Preferred Citation: . The Oceans, Their Physics, Chemistry, and General Biology. New York:  Prentice-Hall,  c1942 1942.


The Oceans Their Physics, Chemistry, and General Biology

H. U. Sverdrup

Professor of Oceanography, University of California
Director, Scripps Institution of Oceanography

Martin W. Johnson

Assistant Professor of Marine Biology, University of California
Scripps Institution of Oceanography

Richard H. Fleming

Assistant Professor of Oceanography, University of California
Scripps Institution of Oceanography
Prentice-Hall, Inc.
New York

Preferred Citation: . The Oceans, Their Physics, Chemistry, and General Biology. New York:  Prentice-Hall,  c1942 1942.

Copyright, 1942, by PRENTICE-HALL, INC. 70 Fifth Avenue, New York




Four years ago when we started the preparation of this book, we hoped to give a survey of well-established oceanographic knowledge, but it soon became apparent that the book could not be brought up to date without summarizing and synthesizing the wealth of information that has been acquired within the past dozen years, as well as the many new ideas that have been advanced. Consequently, the book has grown far beyond its originally planned scope, and the presentation has become colored by the personal concepts of the authors. Discussion of many topics, such as the absorption of radiation in the sea, the relations of organisms to the chemical composition of sea water, or the productivity of the sea, has led to tentative conclusions that are perhaps presented here as better substantiated than is actually the case. At the risk of premature generalizations we have, however, preferred definite statements to mere enumeration o uncorrelated observations and conflicting interpretations, believing that the treatment selected would be more stimulating.

The book is intended to provide a good deal of factual information, but above all it should be an aid to the beginner and specialist alike in the coordination of the various fields of oceanography. The lists of literature at the ends of chapters are not intended to be exhaustive, but will serve as guides to recent publications. When possible, reference is made to books containing comprehensive bibliographies rather than to original papers.

We are much indebted to our colleagues at the Scripps Institution of Oceanography for their numerous helpful suggestions and their constructive criticism of many parts of the book. We are also obliged to Mr. John A. Fleming, Director of the Department of Terrestrial Magnetism, Carnegie Institution of Washington, for permitting free use of unpublished data from the last cruise of the Carnegie.

We extend our thanks to Dr. L. Lek for assistance in carrying out a large number of computations, to Mr. E. C. La Fond for preparation of most of the graphs and charts, to the American Association of Petroleum Geologists and the G. M. Manufacturing Company of New York for use of illustrations published by them, and to the University of Chicago Press for use of the Goode Base Maps. Miss Ruth Ragan has rendered invaluable assistance in correcting and checking manuscript and proofs and in compiling the bibliographies.

The Authors.



Chapter Page
Preface v
I. Introduction 1
II. The Earth and the Ocean Basins
Figure and size of the earth. Distribution of water and land. Relief of the sea floor. Major features of topography. Terminology of submarine topography. Bottom configuration of the oceans. Bottom configuration of the Arctic and Antarctic regions. Bottom configuration of adjacent seas. Submarine canyons. Shorelines. Bibliography.
III. Physical Properties of Sea Water
Salinity and chlorinity. Units of temperature, salinity, and pressure, and their ranges in the sea. Density of sea water. Thermal properties of sea water. Colligative and other properties of sea water. Properties of sea ice. Transmission of sound. Absorption of radiation. Eddy conductivity, diffusivity, and viscosity. Bibliography.
IV. General Distribution of Temperature, Salinity, and Density
The heat budget of the earth as a whole. The heat budget of the oceans. Evaporation from the sea. Salinity and temperature of the surface layer. Theory of the periodic variations of temperature at subsurface depths. Distribution of density. Subsurface distribution of temperature and salinity. The water masses of the oceans. Basins. Bibliography.
V. Theory of Distribution of Variables in the Sea
Scalar fields. Relation between the distribution of properties and the currents in the sea. Distribution of conservative concentrations in the sea. Distribution of nonconservative concentrations. The Principle of dynamic equilibrium. Bibliography.

VI. Chemistry of Sea Water
Constancy of composition. Units used in chemical oceanography. Composition of sea water. Elements present in sea water. Preparation of artificial sea water. Dissolved gases in sea water. The carbon dioxide system. Solubility of salts in sea water. The oxidation-reduction potential of sea water. Inorganic agencies affecting the composition of sea water. Geochemistry of the ocean waters. Bibliography.

VII. Organisms and the Composition of Sea Water
Chemical composition of marine organisms. Interrelations between elements whose distribution is affected by biological activity. Distribution of phosphate, nitrogen compounds, and silicate in the oceans. Factors influencing the distribution of nutrient elements. Compounds of carbon, nitrogen, phosphorus, and silicon in the sea. Bibliography.
VIII. The Sea as a Biological Environment
Physical and chemical characteristics of the marine environment. Other characteristics of the environment. Classification of the marine environment. General character of populations of the primary biotic divisions. Development of life in the sea. Bibliography.
IX. Populations of the Sea
Plant Groups of the Sea: Thallophyta; Blue-green algae (Myxophyceae) ; Green algae (Chlorophyceae) ; Brown algae (Phaeophyceae) ; Red algae (Rhodophyceae) ; Yellow-green algae; The higher plants in the sea. The Animal Population of the Sea: Synopsis of the more important systematic groups of marine animals; Reproduction and life cycles in marine animals. Bibliography.
X. Observations and Collections at Sea
Oceanographic Vessels and Their Facilities: Vessels; Winches; Wire ropes and accessory fittings; Shipboard laboratories. Observations and Collections: Positions at sea; Sonic soundings; Wire soundings; Bottom-sampling devices; Temperature measurements; Water-sampling devices; Treatment and analysis of serial observations; Observations of tides; Deep-sea anchoring. Current Measurements: Drift methods; Flow methods; Current meters; Analysis of records of currents. Collection and Analysis of Biological Samples: Collection of benthic organisms; Collection of nekton; Collection of plankton; Interpretation of plankton observations. Bibliography.
XI. Geneeral Character of Ocean Currents 389
XII. Statics and Kinematics
Statics: Units and dimensions; The fields of gravity, pressure, and mass; Significance of Sigma-T surfaces; Stability. Kinematics: Vector fields; The field of motion and the equation of continuity; Representations of the field of motion in the sea. Bibliography.
XIII. Dynamics of Ocean Currents
The hydrodynamic equations. Currents related to the field of pressure. Friction. Wind currents. Conclusions as to currents on the basis of tonguelike distribution of properties. Thermodynamics of ocean currents. Bibliography.

XIV. Waves and Tides
Introduction. Surface waves. Long waves. Tides. Tidal currents. Effect of friction on tides and tidal currents. The semidiurnal tide of the Atlantic Ocean. Internal waves. Bibliography.

XV. The Water Masses and Currents of the Oceans
Antarctic Circumpolar Ocean. The South Atlantic Ocean. The equatorial region of the Atlantic Ocean. The adjacent seas of the North Atlantic Ocean. The North Atlantic Ocean. Adjacent seas of the Indian Ocean. The Indian Ocean. The South Pacific Ocean. The equatorial region of the Pacific Ocean. The North Pacific Ocean. The adjacent seas of the North Pacific Ocean. The water masses of the oceans: A summary. The deep-water circulation of the oceans. Bibliography.
XVI. Phytoplankton in Relation to Physical-Chemical Properties of the Environment
Methods of Flotation. Factors of Phytoplankton Production: I: Direct primary factors of reproduction and growth; Direct and indirect secondary factors influencing population density. Factors of Phytoplankton Production: 11: Photosynthesis of phyto-plankton; Plant nutrients and vertical circulation of water; Horizontal ocean currents; Temperature. Bibliography.
XVII. Animals in Relation to Physical-Chemical Properties of the Environment
Ecological Groups and Some of Their Adjustments and Conditions of Life: Benthos, animals of the sea floor; Nekton, the swimming animals; Zooplankton, the floating animals. Relations to the Physical-Chemical Properties of the Environment: Light; Salinity; Temperature; Ocean currents; Oxygen. Bibliography.
XVIII. Interrelations of Marine Organisms
Natural Associations of Organisms. Nutritional Relationships: The significance of micro-plants; The significance of micro-animals; Plankton and filter feeders; Detritus feeders and scavengers; Littoral browsers; Preying animals. Biological Factors Influencing Movements and Concentration of Organisms: Phytoplankton-zooplankton; Nekton; Benthos. Marine Bacteria and Their Role in the Biological and Chemical Cycles in the Sea: Structure and reproduction; Bacterial modes of life; The nitrogen cycle ; Phosphorus, carbon, and sulphur cycles; Bacteria and bottom deposits; Distribution of bacteria in the sea. Bibliography.
XIX. Organic Production in the Sea
Phytoplankton production. Zooplankton production. Commercial production. The production in different regions. Bibliography.

XX. Marine Sedimentation
Introduction. Constituents of Marine Sediments. Transportation of Sedimentary Debris: Transportation of sediment to the sea; Transportation of sediment in the sea. Mass Properties of Marine Sediments. Classification of Recent Marine Sediments. Distribution of Pelagic Sediments. Mass Properties of Deep-Sea Sediments. The Environment of Deposition. Calcium Carbonate: Factors which determine accumulation and deposition of calcareous material; The distribution of calcium carbonate. Organic Matter: Quantity and character of organic matter in marine sediments; Distribution of organic matter. Shallow-Water and Nearshore Sediments. Elements Concentrated on the Sea Bottom and Authigenic Minerals. Rates of Sedimentation. Summary of Factors Determining Character of Marine Sediments. Bibliography.

Appendix. Tables for Computing Geopotential Distances Between Isobaric Surfaces
Contents. Symbols and definitions. Explanation. Bibliography.
Index 1061


I. Introduction

Oceanography embraces all studies pertaining to the sea and integrates the knowledge gained in the marine sciences that deal with such subjects as the ocean boundaries and bottom topography, the physics and chemistry of sea water, the types of currents, and the many phases of marine biology. The close interrelation and mutual dependence of the single marine sciences have long been recognized. Thus, the first report, in 1902, of the administration of the International Council for the Study of the Sea, states:

… it was seen from the beginning that the study of the physical conditions, of the chemical nature of the ocean waters, of the currents, etc., was of the greatest importance for the investigation of the problems connected with life, that on the other hand, the study of the floating organisms had particular worth for the solution of hydrographic problems, and consequently that a sharp line should never be drawn between these two main divisions….

The same idea is expressed even more clearly in the prospectus which in 1908 accompanied the first number of the Internationale Revue der gesamten Hydrobiologie und Hydrographie:

Above all, the editors recognize the necessity of a synthesis of our biological and hydrographic-geological knowledge of the waters. These two spheres of investigation are inseparable, since the water, whether as river, lake, or sea, is never a factor in the shaping of the earth without being also a medium for life, and, on the other hand, is never a medium for life without at the same time having an important influence in the shaping of the earth's surface.

As the biology of the waters has now passed from the description of what is found therein to inquire into the causes and origins of the animal and plant life and the phenomena accompanying it, the absolute necessity has arisen for the biologist to really understand the nature of the separate waters, their physics and chemistry as well as their form and the history of their bed.

Since 1900, great advances have been made within all of the marine sciences, and the contacts between the special fields have become more and more intimate. The development is due partly to improved technique and partly to the application to the phenomena in nature of theoretical research and results of laboratory studies.


At the beginning of the twentieth century the distribution of land and sea was known with the exception of parts of the Arctic and the Antarctic, but now most coast lines are charted. The introduction of radio time signals has made possible the exact determination of longitudes and, thus, the correction of minor errors that previously could not be avoided. About twenty-five years ago, knowledge of the submarine topography was very scant, except in shallow-water areas of importance to navigation, where detailed soundings with lead and line could be made rapidly. Since a single sounding in deep water, say 6000 m or more, required several hours, however, such soundings were few and far between, and it was generally considered that the deep-sea bottom was a flat, monotonous plain devoid of rugged relief. In 1911, Fessenden made the first attempts to determine depths by sonic methods, and from about 1920 sonic depth finders have been in use with which soundings can be taken in a few seconds from a vessel running full speed. This new method has in a few years completely altered our concept of the topography of the ocean bottom. Basins and ridges, troughs and peaks have been discovered, and in many areas a bottom topography has been found as rugged as the topography of any mountain landscape.

The increased knowledge of the character of the bottom topography has greatly facilitated the understanding of the flow of the bottom water, and has helped toward explaining observed differences of hydrographic conditions in neighboring areas. Such differences, on the other hand, have been used for determining the height of barriers separating different basins in areas in which few soundings had been made.

Knowledge of bottom sediments has been advanced partly by the introduction of refined physical and chemical methods for the study of the fine-grained inorganic materials and partly by improved methods of obtaining long cores of bottom deposits and specimens of solid rock. The presence of stratification in core specimens from the open ocean has stimulated great interest in the processes controlling the character of the sediments and the rate of sedimentation. These advances have led to the rapid development of the science of submarine geology, which deals with the topographic features of the sea bottom, the agencies that have been active in the development of these features, the types and distribution of sediments, and the processes of sedimentation. The foundation for submarine geology was laid by Sir John Murray, who, with various co-workers, discussed the bottom samples of the Challenger Expedition, 1873–1876, and examined virtually all bottom samples collected prior to his death in 1914. The recent rapid development of interest and the application of new techniques is due mainly to American scientists such as Piggott, Revelle, Shepard, Stetson, and Vaughan, and to the German workers, Correns and Pratje.


Numerous problems within submarine geology cannot be studied properly without knowledge of the nature of ocean currents, the physics and chemistry of the sea water, the general character of the organisms that contribute to the marine sediments, and the transformative activities of the bacteria in the sediments. Thus, submarine geology is dependent upon the results of nearly every other marine science.

Within physical oceanography the study of waves and tides stands in a separate class, because theoretical investigations preceded the accumulation of information as to the exact character of the phenomena. Thus, the theory of surface waves was developed by Gerstner as early as 1802 and was improved by Stokes in 1847. In this early work, however, the water was considered to be an ideal fluid, and many results were more mathematically beautiful than practically applicable, but during the last decades studies, particularly by H. Jeffreys, have partly filled the gap between theory and observation.

Theories of tides were developed by Newton (1642–1727) and Lagrange (1736–1813). Lagrange's formulation of the problems is still valid, but the mathematical difficulties of the theory have not yet been overcome. In recent years notable advances toward the understanding of the tides have been made by the staff of the Liverpool Tidal Institute, headed by Proudman, which has solved fundamental problems for ocean basins of analytically defined geometrical shape, and by the Austrian School, notably Defant and Sterneck, which has used the principles of hydrodynamics in the studies of tides in bays of irregular form. American workers in the U. S. Coast and Geodetic Survey have made notable contributions in the analysis of observed tides and the prediction of tides in coastal areas.

A third type of waves, internal waves or boundary waves, was discussed by Stokes in 1847. The theory treated only the case of waves at the boundary of two ideal fluids of different density, but in 1904 it was successfully applied by Ekman to explain the phenomenon of “dead water.” Subsequent observations indicated that other types of internal waves, generally of tidal periods, were present in the open ocean, and the study of these was greatly furthered by Fjeldstad's theory, which was advanced in 1933 and which deals with internal waves in a liquid whose density increases continuously with depth.

The physical properties of sea water can conveniently be divided into two groups: those that are independent of the ocean currents and of such impurities as suspended particles of inorganic or organic origin—for example, density, specific heat, osmotic pressure, and others; and those that depend upon currents and suspended particles—for example, eddy viscosity, conductivity, diffusion, and transparency. The properties in the first group were accurately determined at the beginning of the century,

although still more refined measurements have been made in recent years. Examination of the properties of the second group is still in rapid development. In the study of the processes of diffusion, knowledge of certain results in marine biology is needed, because distributions of dissolved substances that are influenced by the activity of organisms are often examined. For the explanation of the observed transparency of sea water, biological processes have to be considered, because reduction of the transparency appears in part to be caused by dissolved substances produced by the marine organisms. Knowledge of the physical properties of sea water is essential, on the other hand, in many problems in the other marine sciences.

The general physical theories of ocean currents were developed at the beginning of the present century, notably by Scandinavian oceanographers. Helland-Hansen applied V. Bjerknes' theorem of circulation in a nonhomogenous fluid to the ocean, and Ekman developed the theory of wind-driven currents. The practical application of the Bjerknes theorem was made possible largely through the ingenuity of Fridtjof Nansen, who, about the year 1900, achieved increased accuracy of temperature and salinity determinations. Rossby and his collaborators in the United States have applied results from fluid mechanics to the dynamical problems of the sea. It cannot be foreseen to what extent these new ideas will modify the concepts of the dynamics of ocean currents, but Rossby's work has given new impetus to the theoretical and practical examination of the phenomena. The application of his and of the earlier theories has become increasingly important, owing to the rapid accumulation of temperature and salinity observations and of current measurements.

In the field of physical oceanography, the greater part of the theoretical and practical work can be conducted with little or no attention to results in other marine sciences. Occasionally, conclusions are tested by examining distributions of properties that are influenced by biological activity—for instance, the dissolved oxygen content—but often studies in physical oceanography can be carried out independently. For this reason several oceanographic institutions, such as the Institut für Meereskunde of the University of Berlin, and the Division for Oceanography at the Geophysical Institute, Bergen, Norway, are devoted to research within physical oceanography only, and for this same reason the International Association of Physical Oceanography exists as part of the International Union of Geodesy and Geophysics, and separate from other branches of oceanography. The linking of physical oceanography to the geophysical sciences is logical. Many problems, particularly those related to the dynamics of the atmosphere and the sea, are so similar in meteorology and physical oceanography that the theoretical approach is nearly identical, and the field dealing with the interaction between

the atmosphere and the sea is of equal importance whether considered from the meteorological or the oceanographic point of view.

Although physical oceanography is to a great extent independent of the other marine sciences, its results are used extensively in marine biology and submarine geology. In marine biology it is necessary to know the physical and chemical characteristics of the medium in which the organisms live, the types of currents that may regulate the distribution of the organisms, the vertical motion that carries subsurface waters rich in plant nutrients toward the surface, the depth to which light penetrates, and so on. In submarine geology, knowledge of the large-scale ocean currents is needed for the study of dispersal of fine material brought into the sea by rivers; and information as to the currents at the bottom, their velocity, and their state of turbulence is required for an understanding of the character of sediments found in different localities and of the processes of sedimentation. Because of this wide application it is desirable that the physical oceanographer be acquainted with problems in other marine sciences in order to make his own conclusions better understood and more available to workers in other fields.

In the field of chemistry, the major constituents of the salts that are dissolved in sea water were accurately determined in the 1880's by Dittmar, and around 1900 the empirical relations between chlorinity, salinity, and density were established by Knudsen. These relations are of such importance to physical oceanography that the chemical methods for determining density are considered as necessary tools in that field. Later determinations of the major constituents have introduced only small changes in the early results, but refined methods of analysis have led to the detection of more and more elements in sea water, and in many instances have made possible accurate measurements of the amounts of these minor constituents.

In recent years, notable advances have been achieved in the development of rapid methods for determining the biologically important plant nutrients, and in this field the contacts between marine biology and the chemistry of sea water are so intimate that it is difficult to indicate where the biology stops and the chemistry begins. Concepts and results in physical chemistry have been especially useful in explaining the biologically important carbon dioxide system in the sea water. Important contributions have been made by Harvey, Atkins, and Cooper in Great Britain; Moberg, Rakestraw, and Thompson in the United States; by Wattenberg in Germany, and by Buch in Finland.

Early biological observations were naturally concerned mainly with the larger animals and plants obtainable by simple methods of collecting. This was true even as late as 1839, when Edward Forbes pioneered in the study of marine organisms in relation to their physical environment. It was then generally thought that life could not exist at great depths

because of the adverse living conditions believed to prevail in very deep water. Later this idea of an “azoic zone” below about 600 m was abandoned as a result of positive discoveries of animals inhabiting the abyssal depths.

The realization that life is possible at all depths was of great biological interest, but the discovery of a community of floating microscopic organisms inhabiting the upper water strata was a vastly more significant one, as far as the economy of the sea is concerned. The use of townet and microscope in the study of this ubiquitous multitude of tiny drifting plants and animals of the sea was begun by Johannes Müller about 1846. In 1887 Victor Hensen applied the name plankton to this community and initiated the first quantitative studies. The discovery of plankton made possible an approach to the understanding of the economy of the sea as a whole, for it is the prolific production of this community of organisms that supports the larger forms of life and that is responsible for variations in the distribution of certain chemical elements in the sea.

Most of the older marine biological studies were of necessity chiefly exploratory and descriptive in nature, for only through systematic description and cataloguing of the myriads of forms could a foundation for future work be laid and tools for analysis provided. The descriptive studies must continue, but, as the taxonomic groups become defined and their structure becomes known, the need for such work diminishes and, especially in the better-known areas of the ocean, the emphasis has already shifted to questions of interrelations of the organisms with each other and with the inanimate environment. Near the close of the last century and at the beginning of the present, this point of view was given much impetus by many workers, among whom may be mentioned Brandt, Hjort, Gran, Johnstone, Lohmann, Johannes Schmidt, and Steuer in Europe, and Agassiz, Bigelow, Fraser, Kofoid, and Ritter in America. It is this view which has been carried forward to the present time.

The pressing need for elucidation of the many biological phenomena of the sea has been a motivating force not only in inspiring the application of known aspects of physical and chemical oceanography, but also in stimulating studies of such problems as penetration of light, viscosity, osmotic pressure, the carbon dioxide system, and especially researches on the biologically important elements. Conversely, the distribution and fluctuations of these elements are explicable only through the assistance of biological observations.

During the last decades, for which the trend in oceanographic research has been briefly reviewed, the knowledge of the oceans has been greatly increased. Prior to about 1900, deep-sea observations had been accumulated mainly on large-scale expeditions, foremost among which stands the expedition that represents the beginning of oceanography, the British

Challenger Expedition during the years 1873–1876. Around 1900 the establishment of permanent stations for marine studies gained speed, particularly in northwestern Europe, where the oceanographic investigations were carried out as an assistance to research dealing with fisheries problems of economic importance. Year by year oceanographic institutions and marine biological stations have been added all over the world to the existing stations, and well-equipped expeditions have been sent out, both prior to and after the World War of 1914–1918. Reference to the results from these expeditions will be made repeatedly in the following chapters.

A comprehensive review of the present status of oceanographic exploration and information on the location, history, and facilities of every establishment engaged in marine research was compiled by T. W. Vaughan and published in 1937 by the National Academy of Sciences under the title International Aspects of Oceanography. This book, together with H. B. Bigelow's Oceanography, Its Scope, Problems, and Economic Importance, published by the National Academy in 1931, summarizes the present facilities for oceanographic research and the aims toward which the work is directed. The development of the exploration of the sea is described in such books as Sir John Murray's Summary of the Scientific Results of the “Challenger” Expedition of 1873–1876, Murray and Hjort's Depths of the Ocean, and Herdman's Founders of Oceanography.

Our knowledge of the oceans is still fragmentary and inadequate. In the Pacific and Indian Oceans, large regions exist from which absolutely no information is available, and from most areas only general conditions in certain seasons of the year are known. Expeditions are needed for filling in gaps and for carrying out systematic exploration of regions from which only scattered data are available. The need is even greater for systematic work at sea by well-equipped oceanographic stations that will represent many of the marine sciences, so that findings in different fields can be correlated. Only through such correlation can the marine sciences become the unified science of Oceanography that was visualized at the time the International Council for the Study of the Sea was established.


II. The Earth and the Ocean Basins

Figure and Size of the Earth

As a first approximation the earth may be considered as a sphere, but, according to accurate observations, its figure is more closely represented by an ellipse of rotation—that is, an oblate spheroid, the shorter axis being the axis of rotation. The figure of the earth has been defined by various empirical equations, the constants of which are based on observations and are subject to modification as the number of observations increases and their accuracy is improved. The geometrical figures defined by these equations cannot exactly represent the shape of the earth because of the asymmetrical distribution of the water and land masses.

To define the position of a point on the earth's surface, a system of coordinates is needed, and as such the terms latitude, longitude, and elevation or depth are used. The first two are expressed by angular coordinates and the third is expressed by the vertical distance, stated in suitable linear units, above or below a reference level that is generally closely related to mean sea level. The latitude of any point is the angle between the local plumb line and the equatorial plane. Because the earth can be considered as having the form of a spheroid, and as the plumb line, for all practical purposes, is perpendicular to the surface of the spheroid, any plane parallel to the Equator cuts the surface of the spheroid in a circle, and all points on this circle have the same latitude. These circles are called parallels of latitude. The latitude is measured in degrees, minutes, and seconds north and south of the Equator. The linear distance corresponding to a difference of one degree of latitude would be the same everywhere upon the surface of a sphere, but on the surface of the earth the distance represented by a unit of latitude increases by about 1 per cent between the Equator and the Poles. At the Equator, 1 degree of latitude is equivalent to 110,567.2 m, and at the Poles it is 111,699.3 m. In table 1 are given the percentages of the earth's surface between different parallels of latitude.

The line in which the earth's surface is intersected by a plane normal to the equatorial plane and passing through the axis of rotation is known as a meridian. The angle between two meridian planes through two

points A and B on the surface of the earth is the difference in longitude of the two points. In practice, the longitude is the angular distance measured in degrees, minutes, and seconds east or west from a standard meridian plane, generally that of the Royal Observatory at Greenwich, England. Thus, the longitude is measured from 0° to 180° east and west of Greenwich. On the surface of the earth the linear distance equivalent to unit difference in longitude is nearly proportional to the cosine of the latitude.

Latitude % Cumulative %
0°– 5° 8.68 8.68
5–10 8.62 17.30
10–15 8.48 25.78
15–20 8.30 34.08
20–25 8.04 42.12
25–30 7.72 49.84
30–35 7.36 57.20
35–40 6.92 64.12
40–45 6.44 70.56
45–50 5.92 76.48
50–55 5.33 81.81
55–60 4.71 86.52
60–65 4.05 90.57
65–70 3.36 93.93
70–75 2.64 96.57
75–80 l.90 98.47
80–85 1.15 99.62
85–90 0.38 100.00
SIZE OF THE EARTH (Fowle, 1933)
Equatorial radius, a.................... 6378.388 km
Polar radius, b......................... 6356.912 km
Difference (ab)................ 21.476 km
Area of surface......................... 510,100,934 km2
Volume of geoid......................... 1,083,319,780,000 km3

The distance between points on the earth's surface and the area represented by a given zone cannot be correctly represented unless the size of the earth is known. The values for the equatorial and polar radii are given in table 2, with other data concerning the size of the earth that can be computed from these values. The values for the equatorial and polar radii are those for sea level. The land masses are elevations upon the geometrical figure of the earth, and the sea bottoms represent depressions.

Measurements of depressions below sea level, to be strictly comparable, should be referred to the ideal sea level; that is, to a sea surface which is everywhere normal to the plumb line. In the open ocean the deviations from the ideal sea level rarely exceed 1 or 2 m. The errors that are introduced by referring soundings to the actual sea surface are insignificant in deep water, where the errors of measurement are many times greater. In coastal areas where shoal depths represent a hazard to navigation and where soundings can be made with great accuracy, the

effect of tides is eliminated by referring all soundings in any locality to a single reference plane. Mean sea level is generally used as the plane of reference for elevations on land, and in the Baltic Sea it is employed as the reference level for depths. More commonly, some other plane of reference, indicated on the chart, is used as the reference level for depths; for example:
  1. Mean low water. United States (Atlantic Coast), Argentina, Norway, Sweden.

  2. Mean lower low water. United States (Pacific Coast).

  3. Mean low water springs. Great Britain, Italy, Germany, Denmark, Brazil, Chile.

  4. Mean monthly lowest low water springs. Netherlands.

  5. Lowest low water springs. Brazil, Portugal.

  6. Indian spring low water. India, Argentina, Japan.

  7. Mean semi-annual lowest low water. Netherlands East Indies.

  8. Lowest low water. France, Spain, Norway, Greece.

  9. International low water. Argentina.

The mean of the heights of low-water spring tides is known as the low water springs. International low water is 50 per cent lower, reckoned from mean sea level, than low water springs. Indian spring low water depends upon component tides found by harmonic analysis. Other terms are defined elsewhere (p. 562).

The topographic features of the earth's surface can be shown in their proper relationships only upon globes that closely approximate the actual shape of the earth, but for practical purposes projections that can be printed on flat sheets must be used. It is possible to project a small portion of the earth's surface on a flat plane without appreciable distortion of the relative positions. However, for the oceans or for the surface of the earth as a whole, most types of map projections give a grossly exaggerated representation of the shape and size of certain portions of the earth's surface. The most familiar type of projection is that developed by Mercator, which represents the meridians as straight, parallel lines. Although it is satisfactory for small areas and for the lower latitudes, the size and shape of features in high latitudes are greatly distorted because the linear scale is inversely proportional to the cosine of the latitude. In the presentation of oceanographic materials, this exaggeration is most undesirable and, consequently, projections should be used on which the true shape and size of the earth's features can be more closely approximated.

Numerous types of projections have been developed by cartographers. In some instances, these are geometrical projections of the surface of the geoid on a plane surface that can be flattened out, while in others the essential coordinates, the parallels of latitude, and the meridians have been constructed on certain mathematical principles. Maps and charts

are used for a variety of purposes which require different properties in the projections. Such requirements, incompatible in any single projection, are that the map shall show the features of the earth's surface without distortion of distances, directions, shapes, or areas. The Mercator projection is used for navigational charts because it possesses certain properties that are desirable in navigation. The technique of map making and the properties of the various projections are discussed by Raisz (1938).

In order to show the oceans with the least possible distortion of size and shape, the world maps used in this volume are based on an interrupted projection developed by J. P. Goode. Comparison with a globe will show that the major outlines of the oceans are not distorted and that the margins of the oceans are clearly represented. This projection has the additional advantage of being “equal-area”; that is, that areas scaled from the map are proportional to their true areas on the surface of the earth. To show the relationships between the various parts of the oceans in high latitudes, polar projections are used, and for smaller areas Mercator and other types of projections have been employed.

Distribution of Water and Land

The continental land masses extend in a north-south direction, with the greatest percentage of land in the Northern Hemisphere (table 3), and there is a more or less antipodal arrangement of land- and water-covered areas. The North Polar Sea surrounding the North Pole is opposite to the continent of Antarctica, which is centered on the South Pole, and the continental land masses represented by Europe, Asia, and part of Africa are antipodal to the great oceanic area of the South Pacific. The ocean waters are continuous around Antarctica and extend northward in three large “gulfs” between the continents, on the basis of which three oceans are recognized. The Atlantic Ocean extends from Antarctica northward and includes the North Polar Sea. It is separated from the Pacific Ocean by the line forming the shortest distance from Cape Horn (70°W) to the South Shetland Islands, and the boundary between the Atlantic and the Indian Oceans is placed at the meridian of the Cape of Good Hope (20°E). The boundary between the Pacific and the Indian Oceans follows the line from the Malay Peninsula through Sumatra, Java, Timor, Australia (Cape Londonderry), and Tasmania, and follows the meridian of 147°E to Antarctica. In the north the limit between the Atlantic and the Pacific Oceans is placed in Bering Strait, which is only 58 km wide and has a maximum depth of 55 m. Unless otherwise stated, the oceans as defined above are considered to include the semi-enclosed adjacent seas that connect with them.

Generally speaking, only three oceans are recognized, but it is sometimes desirable to make a further division. The waters surrounding

Antarctica and extending northward about as far as the southern tips of the continents are sometimes designated as the Great Southern Ocean, Austral Ocean, or Antarctic Ocean. In the discussion of the distribution of properties in the sea (chapter XV) it is convenient to define an “Antarctic Ocean” that is limited on the north by a purely oceanographic boundary; namely, the Subtropical Convergence (p. 606).

The nomenclature applied to subdivisions of the oceans is very confused. Generic names designating certain types of features, such as sea, gulf, and bay, are used somewhat indiscriminately and hence have little physiographic significance. For example, the term sea is used in connection with inland salt lakes, such as the Caspian Sea, with relatively isolated bodies of the ocean, such as the Mediterranean Sea, with less isolated areas, such as the Caribbean Sea, and even for some areas with no land boundaries, such as the Sargasso Sea in the western North Atlantic.

Several systems for naming parts of the oceans are employed in oceanographic, work. In certain instances the boundaries are selected arbitrarily by drawing straight or curved lines on the map where there are no land features which constitute natural boundaries. Such a system is followed by the International Hydrographic Bureau (1937). Wüst (1936) has suggested that the submarine ridges that are present at depths of about 4000 m be used to delimit the various parts of the oceans, and that the names now applied to the basins with depths greater than 4000 m be used to designate the areas above them. The general location of such boundaries may be seen in chart I. Oceanography is concerned not only with the form of the oceans as shown on a surface chart, but also with the distribution of properties and living organisms and the nature of the currents. Therefore, a system of nomenclature which indicates the relationships that exist in the sea would be very useful. Wüst's system, based on the ocean bottom topography, meets this purpose for the deep water but not for the upper layers. To formulate “natural regions” of the oceans, other workers, notably Schott (1926, 1935), have attempted to bring together not only geographic and topographic relationships, but also the distribution of properties and organisms, the climatic conditions, and currents. In the discussion of the distribution of organisms, fig. 220 (p. 804) shows how the oceans are subdivided upon the basis of the fauna1 distribution alone, and in the discussion of the water masses of the oceans, fig. 209 (p. 740) shows a subdivision based upon the characteristic temperature and salinity relations of the various regions. A comparison of such charts shows that, although there are certain boundaries which fall in approximately the same localities, there are many regions in which it is not possible to reconcile limits established in different ways.

In table 3 are given the areas of land and water between parallels of latitude five degrees apart. For the whole earth, the ocean waters cover

70.8 per cent of the surface, but the amount of land in the Northern Hemisphere is more than twice that in the Southern Hemisphere, and the water covers only 60.7 per cent of the former and 80.9 per cent of the latter. Between 45°N and 70°N the land exceeds the water surface, whereas between 35°S and 65°S the land forms only a small part of the surface.

Latitude (°) Northern Hemisphere Southern Hemisphere
Water (106 km2) Land (106 km2) Water (%) Land (%) Water (106 km2) Land (106 km2) Water (%) Land (%)
90–85 0.979 ....... 100.0 .... ....... 0.978 .... 100.0
85–80 2.545 0.384 85.2 12.8 ....... 2.929 .... 100.0
80–75 3.742 1.112 77.1 22.9 0.522 4.332 10.7 89.3
75–70 4.414 2.326 65.5 34.5 2.604 4.136 38.6 61.4
70–65 2.456 6.116 28.7 71.3 6.816 1.756 79.5 20.5
65–60 3.123 7.210 31.2 69.8 10.301 0.032 99.7 0.3
60–55 5.399 6.613 45.0 55.0 12.006 0.006 99.9 0.1
55–50 5.529 8.066 40.7 59.3 13.388 0.207 98.5 1.5
50–45 6.612 8.458 43.8 56.2 14.693 0.377 97.5 2.5
45–40 8.411 8.016 51.2 48.8 15.833 0.594 96.4 3.6
40–35 10.029 7.627 56.8 43.2 16.483 1.173 93.4 6.6
35–30 10.806 7.943 57.7 42.3 15.782 2.967 84.2 15.8
30–25 11.747 7.952 59.6 40.4 15.438 4.261 78.4 21.6
25–20 13.354 7.145 65.2 34.8 15.450 5.049 75.4 24.6
20–15 14.981 6.164 70.8 29.2 16.147 4.998 76.4 23.6
15–10 16.553 5.080 76.5 23.5 17.211 4.422 79.6 20.4
10–5 16.628 5.332 75.7 24.3 16.898 5.062 76.9 23.1
5–0 17.387 4.737 78.6 21.4 16.792 5.332 75.9 24.1
90–0° 154.695 100.281 60.7 39.3 206.364 48.611 80.9 19.1
All oceans and seas................................. 361.059 × 106 km2, 70.8%
All land...........:.................................. 148.892 × 106km2, 29.2%

In table 4 are given the areas, volumes, and mean depths of the oceans and of certain mediterranean and marginal seas that together constitute the adjacent seas. The data are from Kossinna (1921), and in most instances the designated areas are readily recognized, but for details concerning the boundaries the original reference should be consulted. The Arctic Mediterranean includes the North Polar Sea, the waters of the Canadian Archipelago, Baffin Bay, and the Norwegian Sea, and is therefore separated from the open Atlantic by the line joining Labrador and Greenland in Davis Strait and running through Greenland, Iceland, Faeroe Islands, Scotland, and England, and across the English

Channel to the Continent. The Asiatic Mediterranean includes the waters eastward of the boundary between the Indian and Pacific Oceans (p. ll), and is separated from the open Pacific by the line extending from the coast of China through Formosa, Philippine Islands, Moluccas, and New Guinea to Cape York, Australia.

Additional data on the areas of the adjacent seas of the Atlantic Ocean are given by Stocks (1938). Littlehales (1932) gives slightly different values for the areas of the oceans.

Relief of the Sea Floor

From the oceanographic point of view the chief interest in the topography of the sea floor is that it forms the lower and lateral boundaries of water. The presence of land barriers or submarine ridges that impede a free flow of water introduces special characteristics in the pattern of circulation and in the distribution of properties and organisms. Furthermore, as will be shown in chapter XX, the nature of the sediments in any area is closely related to the surrounding topography. On the other hand, the geomorphologist or physiographer is concerned primarily with the distribution and dimensions of certain types of topographic features that occur on the submerged portion of the earth's crust. As 71 per cent of the earth's surface is water-covered, knowledge of the major features of the earth's relief will be fragmentary if based only upon those structures that can be seen on land. During the geological history of the earth which covers a span of some thousands of million years, areas now exposed above sea level have at one or more periods been covered by the sea, and parts of the now submerged surface have been above sea level. Many problems in historical geology are therefore dependent upon knowledge concerning the configuration of the sea floor surrounding the continents and the form of the deep-ocean bottom.

Although valuable work in the open ocean has been carried on by scientific organizations, by far the greater proportion of our knowledge of submarine topography is based on soundings taken by or for national agencies in the preparation or improvement of navigational charts. In the United States the U. S. Coast and Geodetic Survey prepares charts for the waters bounding the United States and its possessions, and the Hydrographic Office of the U. S. Navy carries out similar work on the high seas and in foreign waters. The earlier hydrographic work was limited largely to the mapping of coast lines and to soundings in depths less than about 100 fathoms, where hazards to the safe operation of vessels might occur, but deep-sea soundings received a great impetus when surveys were made prior to the laying of the transoceanic cables in the latter part of the nineteenth century. Up to and including the time of the voyage of the Challenger, 1873–1876, all soundings were made with hemp ropes, which made the process a long and tedious undertaking,

since in depths of several thousand meters it required several hours to make a single depth measurement. A great improvement in sounding equipment was made when wire ropes and, for sounding at great depths, single-strand wire were introduced, thereby reducing the size of the gear necessary for deep-sea sounding and the time for lowering and reeling in. An earlier improvement was achieved by the introduction of methods of dropping the sounding weight after it had reached the bottom, thus reducing the load to be reeled in. The first such device was developed in the middle of the last century by J. M. Brooke of the U. S. Navy.
Equipment now used for wire sounding is described in chapter X, which deals with oceanographic apparatus and methods.

Body Area (106 km2) Volume (106 km3) Mean depth (m)
Atlantic Ocean excluding adjacent seas 82.441 323.613 3926
Pacific Ocean excluding adjacent seas 165.246 707.555 4282
Indian Ocean excluding adjacent seas 73.443 291.030 3963
All oceans (excluding adjacent seas) 321.130 1322.198 4117
Aretic Mediterranean 14.090 16.980 1205
American Mediterranean 4.319 9.573 2216
Mediterranean Sea and Black Sea 2.966 4.238 1429
Asiatic Mediterranean 8.143 9.873 1212
Large mediterranean seas 29.518 40.664 1378
Baltic Sea 0.422 0.023 55
Hudson Bay 1.232 0.158 128
Red Sea 0.438 0.215 491
Persian Gulf 0.239 0.006 25
Small mediterranean seas 2.331 0.402 172
All mediterranean seas 31.849 41.066 1289
North Sea 0.575 0.054 94
English Channel 0.075 0.004 54
Irish Sea 0.103 0.006 60
Gulf of St. Lawrence 0.238 0.030 127
Andaman Sea 0.798 0.694 870
Bering Sea 2.268 3.259 1437
Okhotsk Sea 1.528 1.279 838
Japan Sea 1.008 1.361 1350
East China Sea 1.249 0.235 188
Gulf of California 0.162 0.132 813
Bass Strait 0.075 0.005 70
Marginal seas 8.079 7.059 874
All adjacent seas 39.928 48.125 1205
Atlantic Ocean 106.463 354.679 3332
Pacific Ocean, including adjacent seas 179.679 723.699 4028
Indian Ocean 74.917 291.945 3897
All oceans (including adjacent seas) 361.059 1370.323 3795

Because of their practical importance and the ease with which they could be obtained, the number of soundings in depths less than a few hundred meters accumulated rapidly during the nineteenth century, but in 1895 there existed only 7000 soundings from depths greater than about 2000 m, and of these only about 550 were from depths greater than 5500 m (Bencker, 1930). These data were used by Murray in preparing the bathymetric charts accompanying the reports of the Challenger Expedition.

During the next twenty-five years the number of deep-sea soundings increased slowly, but the introduction of sonic-sounding equipment after 1920 has completely changed the picture. Devices for measuring the depth by timing the interval for a sound impulse to travel to the sea bottom and back again (only a few seconds even in deep water) are used in surveying work and are now standard equipment on many coastwise and oceanic vessels. The development of automatic echo-sounding devices (chapter X) not only made depth measurements simple but, by making accurate bathymetric charts available, introduced another aid in navigation, since passage over irregularities of the sea floor may be used to check positions. This development has necessitated extending accurate surveys into deeper water and, hence, farther from shore. Along the coasts of the United States the bottom is now being charted in detail to depths of about 4000 m. With sonic methods, if the appropriate apparatus is available, it is no more trouble to sound in great depths than it is in shoal waters, and, since many naval vessels and transoceanic commercial vessels make systematic records of their observations, the soundings in the deep sea are now accumulating more rapidly than they can be plotted.

The most common method of representing submarine topography is to enter upon a chart showing the coast lines the numerical values of the soundings at the localities in which they were obtained. Charts issued by the national hydrographic services of the English-speaking countries give depths in fathoms or, if harbor charts, in feet (1 fathom = 6 ft = 1.8288 m). Those issued by other countries generally use meters, although still other units are employed by certain European countries.

Because it is generally impossible to enter all soundings, and as numerical values alone do not give any graphic representation of the topography, contours of equal depths (isobaths) are drawn in those regions in which the number of soundings or the purpose of the chart makes it desirable. On navigational charts, contours are generally restricted to shallow areas where soundings are also shown, but, for certain regions that have been carefully examined, charts are now issued with contours entered to depths as great as 2000 fathoms (for example, U. S.

Coast and Geodetic Survey Chart no. 5101A for the southern California coast, issued in 1939). In bathymetric charts prepared for oceanographic work, contours alone may be shown, but only in exceptional cases, such as in physiographic studies, are the isobaths shown for equal intervals of depth.

The accuracy with which submarine topography can be portrayed depends upon the number of soundings available and upon the accuracy with which the positions of the soundings were determined. Topographic maps of land surfaces are based on essentially similar data; namely, elevations of accurately located points, but the surveyor has one great advantage over the hydrographer. The surveyor is able to see the area under examination and thereby distribute his observation points in such a manner that the more essential features of the topography are accurately portrayed. The hydrographer, on the other hand, must construct the topography of the sea floor from a number of more or less random soundings. Sonic sounding methods and the introduction of more accurate means of locating positions at sea (see Veatch and Smith, 1939) have made it feasible to obtain adequate data for constructing moderately accurate charts or models of parts of the sea floor. This is particularly true of the coastal waters of the United States. Veatch and Smith have prepared contour maps of the eastern seaboard based on the investigations of the U. S. Coast and Geodetic Survey, and Shepard and Emery (1941) have made use of similar data from the Pacific Coast, where over 1,300,000 soundings were available.

In some instances it is preferable to represent the bottom configuration by vertical profiles or by relief models, but, because of the difference in magnitude of the vertical and horizontal dimensions of the oceans, it is generally necessary to exaggerate the vertical scale. The average depth of the ocean is about 3800 m, and the vertical relief of the ocean floor is therefore of the order of a few kilometers, whereas the horizontal distances may be of the order of thousands of kilometers. Hence such distorted representations give a false impression of the steepness of submarine slopes. If profiles are drawn to natural scale, the ocean waters form a shallow band with barely perceptible undulations of the bottom. Examples of undistorted profiles are given by Johnstone (1928).

In fig. 1 are shown two representations of a profile of the sea bottom in the South Atlantic based on the observations of the Meteor (Stocks and Wüst, 1935). The upper section (A) is constructed from thirteen wire soundings, and is comparable in detail to most of the profiles that could be prepared before the introduction of sonic methods. The lower section (B) is based upon over 1300 sonic soundings that were taken by the Meteor along the same route, shown in the map at the bottom part of the figure (C), where the depth contours are from chart I. The increasing complexity of the known topography of the sea bottom resulting

from the greater number of observations is readily seen from a comparison of A and B.

The water surface coincides, for all practical purposes, with the surface of the geoid, and the sea bottom, if “flat,” would be parallel to the sea surface. Irregularities of the sea floor therefore represent departures from this surface, which is convex outward. Only in small features with steep slopes are depressions actually concave outward.


Bottom topography in the South Atlantic Ocean. (A) Profile of the bottom between the South Shetland Islands and Bouvet Island based on 13 wire soundings. (B) Profile over the same course constructed from over 1300 sonic soundings (Meteor). (C) Bottom configuration as shown in Chart I and the track of the Meteor. Vertical exaggeration in (A) and (B) about 200:1. (In part, after Stocks and Wüst, 1935.)

The greatest depths so far discovered are in the Pacific Ocean, where, in the Philippines Trench and the Japan Trench, soundings greater than 10,000 m have been obtained. In the Philippines Trench the German vessel Emden obtained a sonic sounding of 10,540 m, which, however, is considered to be about 200 m too great. The Dutch vessel Willebrord

Snellius, working in the same area and using wire sounding equipment, found depths greater than 10,000 m. A sounding of 10,550 m in the Japan Trench was made by the U.S.S. Ramapo, using sonic equipment. The deepest sonic sounding in the Atlantic Ocean, 8750 m, was obtained in the Puerto Rico Trough by the U.S.S. Milwaukee. Bencker (1930) has listed the numerous deeps as well as the shoals found in oceanic areas.

Representations of submarine topography are usually referred to sea level, and particular interest has always been attached to those regions in which great depths are found. The greater detail with which the sea floor can now be mapped has emphasized the importance of relative relief; that is, the form and magnitude of elevations or depressions with respect to their general surroundings. In later pages it will be shown that there are two primary levels of reference on the earth's crust, one slightly above sea level, corresponding to the land masses, and a second at depths between 4000 and 5000 m, corresponding to the great oceanic basins. In comparing topographic features on land with those on the sea floor it is essential to consider them with reference to these levels.


Hypsographic curve showing the area of the earth's solid surface above any given level of elevation or depth. At the left in the figure is the frequency distribution of elevations and depths for 1000-meter intervals.

One method of presenting the character of the relief of the earth's crust is by means of a hypsographic curve showing the area of the earth's solid surface above any given contour of elevation or depth. The hypsographic curve in fig. 2 is from Kossinna (1921). Although added data

concerning bottom configuration may slightly modify this curve, the general features will not be changed. The high mountains form a relatively small part of the land surface, and hence the mean elevation of the subaerial crust is only 840 m. The large areas of low-lying land have their counterpart in the relatively large areas in shallow water between the surface and approximately 200 m (table 5). These coastal areas of shallow depth correspond to the continental shelves. Below the continental shelf there is a relatively small area of depths between 200 m and 3000 m, corresponding to the continental slope, and then follows the extensive oceanic abyss, with depths between about 3500 and 6000 m. The deeps, which by definition exceed 6000 m, form a very small part of the sea floor. Shown in the figure is the mean sphere depth, which is the uniform depth to which the water would cover the earth if the solid surface were smoothed off and were parallel to the surface of the geoid. The mean depth of the sea, which is 3800 m, is also shown.

The hypsographic curve of the earth's crust should not be interpreted as an average profile of the land surface and sea bottom, because it represents merely the summation of areas between certain levels without respect to their location or to the relation of elevations and depressions. Actually, the highest mountains are commonly near the continental coasts, large areas of low-lying land are located in the central parts of the continents, and the greatest depths are found near the continental masses, and not in the middle of the oceanic depressions. Entered in fig. 2 are the percentages of elevations and depressions for 1000-m intervals. These show two maxima, one just above sea level and a second between depths of 4000 and 5000 m. The significance of these maxima is discussed later (p. 23).

In table 5 are given the percentage areas of the depth zones in the three oceans, and for all oceans with and without adjacent seas. It will be noted that the shelf (0–200 m) represents a prominent feature in the Atlantic Ocean, which is also the shallowest of the oceans. By combining data in tables 4 and 5 the absolute areas of the depth zones may be computed. The hypsographic curve in fig. 2 is based on the values for all oceans, including adjacent seas.

During the geological history of the earth, great changes have occurred in the relief of the land and sea bottom. The exact nature and extent of these vertical movements is beyond the scope of the present discussion, but it should be noted that changes in relative sea level of the order of 100 m, which are readily accounted for by the withdrawal and addition of water during glacial and interglacial periods, would expose and inundate relatively large areas.

The continental shelf is generally considered to extend to depths of 100 fathoms, or 200 m, but Shepard (1939) found that the limit should be somewhat less than this; namely, between 60 and 80 fathoms (110 and

146 m). Profiles of the shelf show the presence of many minor terraces that may represent the effects of waves and currents when the sea level stood at a lower level, probably during the glacial periods. Veatch and Smith (1939), from their detailed study of the continental shelf off the Atlantic coast of the United States, found many small ridges approximately parallel to the coast line. The continental shelf varies greatly in width and slope. In some cases, as off mountainous coasts, the shelf may be virtually absent, whereas, off glaciated coasts and off the mouths of large rivers and areas with broad lowlands, the shelf may be very wide. For the world as a whole, the shelf width is approximately 30 miles, with a range from zero to over 800 miles. This extremely wide shelf is found in the North Polar Sea along the coast of Siberia.

Depth interval (m) Including adjacent seas Excluding adjacent seas
Atlantic Pacific Indian All oceans Atlantic Pacific Indian All oceans
0–200 13.3 5.7 4.2 7.6 5.6 1.7 3.2 3.1
200–1000 7.1 3.1 3.1 4.3 4.0 2.2 2.7 2.8
1000–2000 5.3 3.9 3.4 4.2 3.6 3.4 3.1 3.4
2000–3000 8.8 5.2 7.4 6.8 7.6 5.0 7.4 6.2
3000–4000 18.5 18.5 24.0 19.6 19.4 19.1 24.4 20.4
4000–5000 25.8 35.2 38.1 33.0 32.4 37.7 38.9 36.6
5000–6000 20.6 26.6 19.4 23.3 26.6 28.8 19.9 26.2
6000–7000 0.6 1.6 0.4 1.1 0.8 1.8 0.4 1.2
>7000 .... 0.2 .... 0.1 .... 0.3 .... 0.1

From the above values it may be seen that the average slope of the shelf is of the order of 2 fathoms per mile, or 0.2 per cent. This corresponds to a slope angle of about 7ʹ. Although there is a general seaward slope of the shelf, it is by no means an even-graded profile. As mentioned above, there may be terraces, ridges, hills, and depressions, and in many areas there are steep-walled canyons cutting across it. Shelf irregularities are most conspicuous off glaciated coasts, and were caused by the ice during a glacial period when this zone was exposed to glacial erosion (Shepard, 1931).

On land the slope is often more significant than the absolute range in elevation. According to Littlehales (1932) the smallest slope that the human eye can detect is 17ʹ. Therefore, except for the minor irregularities, the continental shelf would in general appear flat.


From an examination of 500 profiles, Shepard (1941) found that the inclination of the continental slope varied with the character of the coast. Continental slopes off mountainous coasts have, on the average, a slope of about 6 per cent (3°30ʹ), whereas off coasts with wide, well-drained coastal plains the slopes are about 3.5 per cent (2°0ʹ).

The submerged slopes of volcanic islands are similar to the exposed slopes of volcanic mountains, and may be as great as 50° (Kuenen, 1935). In large submarine canyons the walls are as rugged and precipitous as those of the Grand Canyon of Arizona (fig. 8, p. 40). Fault scarps above and below sea level show comparable slopes.

The average slopes of the deep-sea floor are small. Krümmel (Littlehales, 1932) found that in the North Atlantic the mean slopes varied between about 20ʹ and 40ʹ, but these are averages, or were obtained by dividing the difference in elevation by the distance between two points. Where the distances are great or when the number of soundings is small, the slopes obtained in this way do not give a true representation of the relief. The increased data now available have revealed irregularities comparable in ruggedness to the larger topographic features on land.

Major Features of Topography

The discussion of the bottom topography of the oceans will be restricted to a brief consideration of the large-scale topographic features that are represented on small charts with large contour intervals. In regions where many soundings have been obtained, it has been found that the sea bottom may be virtually as irregular as the land surface, but such details can be shown only on large-scale charts with small contour intervals, and are not included in this volume.

Submarine geology is concerned with the topography of the sea floor, the composition and physical character of the sedimentary and igneous materials that are found on the ocean bottom, and the processes involved in the development of topographic relief. The field is a relatively new one which has received great impetus from the development of sonic sounding methods that made it possible to obtain accurate maps of the sea floor, and from the development of geophysical methods (measurement and interpretation of gravity anomalies, of the earth's magnetic field, and of the path and velocity of earthquake and artificial seismic waves) that yielded estimates of the character and thickness of the materials forming the crust of the earth. However, there is yet no agreement concerning the processes involved in the geological history of the ocean basins, and the various hypotheses will not be discussed here. General reviews of the problems will be found in Johnstone (1928), Bucher (1933), Kuenen (1935), and Gutenberg (1939). A symposium on the geophysical exploration of the sea bottom (Field et al, 1938) covers many of the developments.


The distribution of elevations and depressions on the earth's crust (fig. 2) shows that there are large portions with elevations between sea level and 1000 m, and with depths between 4000 m and 5000 m. According to Bucher (1933) the larger, lower ones are related to the character of the earth's crust, while the upper ones are the result of subaerial erosion and sedimentation. The question then arises as to the extent to which the topography of the ocean bottom with reference to a depth of about 4500 m corresponds to that of the land with reference to sea level or a slightly higher level. According to Bucher, the large-scale features are essentially similar, and elevations and depressions of comparable dimensions are found both on land and on the ocean bottom. Although the major features are comparable, the details are quite different, because erosion, which plays such an all-important role in the creation of sharp relief and in the ultimate destruction of elevations on land, is virtually absent in the sea. In the sea the most effective agents of erosion are the surface waves, and these tend to produce flat-topped features that are restricted to shallow depths, since the velocity of the water particles in such waves decreases rapidly with increasing depth (p. 528). Other processes which may contribute to erosion of the sea floor are discussed in chapter XX and in the section dealing with the origin of submarine canyons (p. 41). Deposition is the characteristic process that modifies the topography of the sea bottom. Sedimentary debris accumulates in depressions, while there is little or no accumulation on topographic highs, which are devoid of fine-grained sediment and are subject to erosion if near the surface or in localities of exceptionally strong currents.

Bucher (1933) has stated that there are essentially two types of large-scale topographic features on the land and on the ocean bottom: (1) those of approximately equidimensional lateral extent, to which he applies the names swells and basins, and (2) those of elongate form, generally with steeper sides, to which he applies the names welts and furrows. On the ocean bottom the elongate welts and furrows appear to be the more common, and there is a considerable range in the size of such structures. There is a tendency for the large welts on the sea bottom to be parallel to the continental coasts, so that the oceans are divided into elongate troughs. Transverse ridges in turn subdivide these major depressions into a series of basins that are separated from one another to a greater or lesser degree. This ridge and basin topography is clearly shown by the bottom of the Atlantic Ocean and the Indian Ocean and in the western part of the Pacific Ocean, but does not appear to be so conspicuous a feature in the main part of the Pacific.

Within the smaller welts and furrows, the steepest slopes, the highest elevations, and the greatest depths are found. The welts and furrows are commonly close together, with arced outlines, and are characteristically found near the continents. The deep furrows are generally on the convex

side of the arc-like welts, but in some instances great depths are found on the concave side of the ridges (chart I). Parts of the welts may extend above sea level, forming an island or chain of islands. Such features are found in the northern and western parts of the Pacific Ocean, in the East Indian Archipelago, in the West Indies, and in the region between South America and Antarctica. Along the west coast of South America the welt corresponds to the mountain chain of the Andes, and is therefore part of the continent. That these structures are of relatively recent origin and are in regions of crustal instability is shown by the presence of extinct or active subaerial or submarine volcanoes on the welts and by the fact that they are in regions of the most pronounced gravity anomalies and are the sites of great seismic activity (Field et al, 1938).

Terminology of Submarine Topography

The terms applied to features of submarine topography will be classified according to the origin of the features rather than according to their size, although the latter procedure is the common one (for example, Niblack, 1928, Littlehales, 1932). The features of submarine relief may be grouped in two main categories, depending upon whether they have gained their characteristic form through diastrophic activity (crustal movements) or through erosion or deposition. The primary large-scale process involved in the development of relief must be diastrophic, but in many cases the characteristic feature is produced by erosion or deposition. No distinction will be made here between features that have been formed below the sea surface and those that may possibly owe their origin to subaerial erosion or deposition. As pointed out before, deposition in the sea tends to fill in the depressions and thus to level out the minor irregularities of the bottom, and, with the exception of those cases in which organisms play an important role (for example, in the formation of coral reefs), little or no deposition takes place on topographic highs.

There has been much discussion as to the processes that have led to the formation of the continental and insular shelves. Some authors maintain that they are wave-built (depositional); others consider that they are wave-cut (erosional), or that they are a combination of both processes (Johnson, 1919; Shepard, 1939). Geophysical studies on the two sides of the North Atlantic (Bucher, 1940) indicate that the shelves are composed of great prism-shaped accumulations of sedimentary rock that at the outer edge of the shelf bordering the eastern United States are 4000 m thick. To what extent these features resulted from the slow accumulation and sinking of the crust and to what extent violent diastrophic movements have been involved has not yet been decided. The characteristic form of the shelf and of isolated flat-topped banks and shoals, and other features of the shallow bottom indicate that wave erosion and transportation by currents have played an important part

in their development. Particular importance is attached to the relative lowering and rise of sea level that took place through the accumulation of continental ice during each glacial period and the subsequent melting. According to Daly (1934) the maximum lowering of sea level was of the order of 100 m, but recently Shepard (Shepard and Emery, 1941), in order to account for the origin of the submarine canyons, has advanced arguments in favor of a lowering of the order of 1000 m. Even a lowering of 100 m would expose large areas of the shelf and would explain the presence of the submerged terraces and other irregularities which could have been produced by wave action when the water stood at a lower level.

The terms used to designate certain types of topographic features, their French and German equivalents, and their definitions, which are given below, correspond to those suggested by the International Hydrographic Bureau (Niblack, 1928). Unfortunately, there is still considerable confusion in the use of certain terms, particularly those which apply to the larger features of the topography. Sometimes several different descriptive terms have been applied to the same structure, and in other instances the same term is applied to features of vastly different size and probably of different origin. A committee of the International Association of Physical Oceanography (Vaughan et al, 1940) attempted to clarify many of the problems relating to the terminology, but much confusion still prevails. In order to designate any individual feature, the descriptive term is prefixed by a specific name. The specific names attached to large-scale features are generally geographical, but those assigned to such features as banks, shoals, seamounts, canyons, and sometimes deeps are often those of vessels or individuals associated with their discovery or mapping.

Features Resulting from Crustal Deformation

  1. Elevations. The large-scale elevations of the ocean bottom are termed ridges, rises, or swells.

    1. Ridge (F, Dorsale; G, Rücken). A long and narrow elevation with sides steeper than those of a rise.

    2. Rise (F, Seuil; G, Schwelle). A long and broad elevation which rises gently from the ocean bottom.

      Isolated mountain-like structures rising from the ocean bottom are known as seamounts. Where the ridges are curved, and particularly if parts of them rise above sea level, they are sometimes termed arcs. The broad top of a rise is termed a plateau. The expression sill is applied to a submerged elevation separating two basins. The sill depth is the greatest depth at which there is free, horizontal communication between the basins.

  2. 26
  3. Depressions. The terms trough, trench, and basin are those most commonly applied to the large-scale depressions on the ocean bottom.

    1. Trough (F, Dépression; G, Mulde). A long, broad depression with gently sloping sides.

    2. Trench (F, Fossé; G, Graben). A long and narrow depression with relatively steep sides.

    3. Basin (F, Bassin; G, Becken). A large depression of more or less circular or oval form.

      The terms defined above are used rather loosely and are applied to features of a wide range in size.

      For those parts of a depression which exceed 6000 m in depth, the term deep (F, Fosse; G, Tief) is used. As originally suggested by Murray, the term designated areas where the depths exceeded 3000 fathoms (5486 m), but it is now generally restricted to those depressions of greater depth (Vaughan et al, 1940). The term depth (F, Profondeur; G, Tiefe), prefixed by the name of the vessel concerned, may be used to designate the greatest sounding obtained in any given deep.

Features Resulting from Erosion, Deposition, and Biological Activity

As pointed out above, the features in this category have been produced by erosion of, or deposition upon, structures which may be primarily of diastrophic origin. The most prominent types of features in this group are the shelf and the slope.

  1. Shelf. The zone extending from the line of permanent immersion to the depth, usually about 120 m, where there is a marked or rather steep descent toward the great depths. Continental Shelf (F, Plateau continental; G, Kontinental-Schelff) is applied to the feature bordering the continents, while Insular Shelf (F, Socle; G, Insel-schelff) is used for the feature surrounding islands.

  2. Slope. The declivity from the outer edge of the shelf into deeper water. Continental Slope (F, Talus continental; G, Kontinental-Abfall) and Insular Slope (F, Talus insulaire; G, Inselabfall) are applied to the slopes bordering continents or islands.

The following terms are applied to the upper parts of elevations which show the effects of erosion or deposition.

  1. Bank (F, Banc; G, Bank). A more or less flat-topped elevation over which the depth of water is relatively small, but which is sufficient for surface navigation.

  2. 27
  3. Shoal (F, Haut-fond; G, Untiefe or Sandgrund). A detached elevation with such depths that it is a danger to surface navigation and which is not composed of rock or coral.

  4. Reef (F, Récif; G, Riff). A rocky or coral elevation (generally elongate) which is dangerous to surface navigation. It may extend above the surface.

    A variety of names has been applied to the steep-walled fissures that penetrate the slope and cut across the shelf. The most commonly used terms are canyon and valley, but gully, gorge, and mock-valley are also applied to these features.

In addition to the terms given above, many expressions are employed in descriptions of submarine topography with the same meanings that they have when used for land topography.

Bottom Configuration of the Oceans

The major features of the topography of the ocean bottom are of such large dimensions that they are readily shown on a chart with contour intervals 1000 m apart. Such a representation is given in chart I, where the contours are entered for 1000-m intervals between 3000 m and 7000 m. The areas with depths less than 3000 m represent a rather small part of the sea floor, and the complex nature of the contours for depths less than this would confuse rather than add to the value of a chart of this kind. The topography is based upon the most recent charts available, and primarily upon the bathymetric chart prepared by the International Hydrographic Bureau in 1939 (Vaughan et al, 1940). Other sources that may be consulted for details concerning the configuration of the ocean floor are listed on page 29. It will be noted that the complexity of the topography varies in different regions. This difference must be attributed, in part, at least, to the variable amount of data available, because in those regions where the soundings are widely spaced the contours will be smooth and rounded, whereas in those areas where there are numerous soundings the contours are more complex and irregular. The Atlantic Ocean, the central part of the North Pacific Ocean, the Northern Indian Ocean, and the area surrounding Antarctica are fairly well sounded, but in many other regions, such as the North Polar Sea and the Southern Indian and South Pacific Oceans, the observations are very sparse. The increase in the complexity of the known topographic features that follows the accumulation of more depth measurements can be seen by comparing recent bathymetric charts with those published in the early years of the present century. The status of bathymetric knowledge in 1937 is shown by a series of charts in Vaughan et al (1937).

As stated above, the topography of the ocean bottom is characterized by depressions and elongated ridges. Some of these features are of very

great dimensions, as may readily be seen from chart I. Longitudinal ridges divide the three oceans into elongated troughs. This feature is most conspicuous in the Atlantic Ocean, where the Atlantic Ridge, extending from Iceland to Bouvet Island in about 55°S, separates the western and the eastern troughs. Depths exceeding 5000 m exist on both sides of the ridge, which is continuous at depths less than 3000 m over the greater part of its length and in several places extends above sea level. There is one small but significant break in the ridge, in the Romanche Furrow just north of the Equator, where the saddle depth is located between 4500 and 4800 m. The Walfisch Ridge, which extends northeast from the Atlantic Ridge in the vicinity of Tristan da Cunha (37°S) to the coast of Africa in latitude 20°S, is continuous at 3500 m and almost so at 3000 m. The Rio Grande Ridge extends westward from the Atlantic Ridge (30°–35°S) and is almost continuous at 4000 m. The presence of these two transverse ridges has a pronounced effect on the deep-water circulation in the Western and Eastern Atlantic and, hence, on the distribution of properties (chap. XV).

A longitudinal ridge, the Indian Ridge, is present in the Indian Ocean and extends from India to Antarctica, but differs from the one in the Atlantic Ocean in that it is wider and does not extend so near the surface. In the Pacific Ocean the longitudinal elevations are not so conspicuous; however, the West Pacific Ridge, which is actually composed of several shorter ridges, can be traced from Japan to Antarctica, and is continuous at depths less than 4000 m except for breaks at 11°N, 10°S, and 53°S. A second elevation extends from Central America to the south and west, reaching Antarctica in the longitude of New Zealand. This East Pacific Ridge is continuous at depths less than 4000 m and separates the central depression from the deep basins bordering Central and South America and the Pacific Antarctic Basin. The effect of these major elevations on the distribution of bottom-water temperatures is shown in fig. 211, p. 749.

Within the major depressions or troughs which are bordered by the continents and the longitudinal ridges are transverse ridges that separate to a greater or lesser degree a number of basins. Wüst (Vaughan et al, 1940) has suggested that the 4000-m contour be used as the boundary in designating basins, but this is a purely arbitrary delimitation that places undue emphasis upon the absolute depth rather than upon the relative relief, which in many instances is of greater significance. For example, the Mediterranean Sea Basin is virtually excluded from such a classification, although it is a deep, isolated basin, much of it extending more than 3000 m below the sill in the Strait of Gibraltar. In the tabulation accompanying chart I are listed the names for the major parts of the oceanic depressions which Wüst has termed basins; namely, those parts which have depths exceeding 4000 m. Certain individual basins are clearly

defined by the presence of ridges that are continuous at 4000 m, but names have also been applied to various parts of troughs, in which case the boundaries are located at the shallowest or narrowest part of the depression. In some areas the nomenclature is incomplete, and a single name is applied to a number of more or less isolated depressions, such as the Madagascar Basins. In the Central Pacific, where the knowledge of the topography has increased rapidly in recent years and where the basin and ridge type of topography does not appear to be present, no names have been applied. The names used to designate the major features of the topography are discussed at length by Vaughan et al (1940), and the names indicated on chart I generally conform to the recommendations made in their report.

In the tabulation of the basins given on chart I are listed some of the more prominent deeps; namely, those features where the depths exceed 6000 m. Some deeps are located more or less centrally in the large basins; for example, Wharton Deep, Byrd Deep, and the numerous deeps in the central part of the North Pacific, but these rarely exceed 7000 m in depth. On the other hand, numerous deeps of elongate character are located near and parallel to continental coasts, island arcs, or submarine ridges which correspond to the furrows discussed on p. 23. These marginal deeps, to which the term trench or sometimes trough is applied, are the features within which the greatest depths are found, in nearly all cases exceeding 8000 m. Only one such trench is found in the Indian Ocean; namely, the Sunda Trench. In the Atlantic are to be found the Romanche Trench, the South Sandwich Trench, and the Puerto Rico and Cayman Troughs. The greatest number are in the western part of the Pacific Ocean, although there is a chain of such features paralleling the mountainous coast of parts of Central and South America. As stated before, the regions in which these deep trenches occur are sites of volcanic and seismic activity. The complex topography of the East Indian Archipelago, which has been described by Kuenen (1935), is shown schematically in fig. 208, p. 736.

For a detailed description of the features of the sea bottom the reader should consult Littlehales (1932). Vaughan (1938) has described the topography of the Southern Hemisphere. There is much information of value in the large report by Vaughan and others (1940), which also contains the small-scale bathymetric chart prepared by the International Hydrographic Bureau on a Mercator projection, a special chart of the North Pacific prepared by the U. S. Hydrographic Office, and an excellent, detailed chart of the Caribbean Sea region prepared by the same agency. The standard charts on the bathymetry of the oceans are those in the series known as the Carte Générale Bathymétrique des Océans, published by the International Hydrographic Bureau at Monaco. These charts comprise twenty-four sheets which are revised from time to time

and are issued periodically as the new charts are completed. The third edition is now being issued. On these charts the depths are entered in meters. General bathymetric charts of the oceans are included in the publications of Schott (1926, 1935). Detailed charts of restricted regions and general charts of the oceans are issued by the various national agencies responsible for the publication of navigational charts. The reports of the Meteor Expedition (Deutsche Atlantische Expedition “Meteor” 1925–1957, Wissenschaftliche Ergebnisse) will contain charts of the Atlantic Ocean, to be issued in thirteen sheets. The Snellius Expedition has produced excellent bathymetric charts of the East Indian region (van Riel, 1934). The Geological Society of America has sponsored the preparation and publication of detailed topographic charts of the eastern and western coasts of the United States based upon the soundings made by the U. S. Coast and Geodetic Survey (Veatch and Smith, 1939; Shepard and Emery, 1941).


Polar projection of the Arctic regions showing the generalized topography of the sea bottom. (Cherevichny's soundings of 1941 not included.)

Bottom Configuration of the Arctic and Antarctic Regions

Fig. 3 has been prepared to show the submarine topography of the North Polar regions, which cannot be properly visualized from the interrupted projection used in chart I. The figure is based on a chart by Stocks (1938) and incorporates all of the available data. Because of the larger scale it is possible to show the contours for the shallower depths that form such a large part of this area. The conspicuous topographic

features are the deep, partially isolated basins and the very wide shelf from which rise the large islands of the Canadian Archipelago, Greenland, and most of the islands to the north of Europe and Asia.

Very little is known of the topography of the North Polar Basin, and the form of the contours is largely hypothetical. Soundings greater than 3000 m are fairly numerous to the north of Europe, and there are some to the north of Alaska. A line of soundings also extends from the Pole and parallels the east coast of Greenland. These soundings were obtained by the Russian expedition which landed on the ice from planes and, in 1937–1938, drifted with the pack ice until picked up off the east coast of Greenland. Within 100 km of the Pole, this party obtained a sounding of 4300 m. The 5000-m contour is inserted on the basis of a single sounding of 5440 m obtained in 1927 by Sir Hubert Wilkins, who flew out by plane from Alaska, landed on the ice, and measured the depth with a portable sonic sounding instrument. The correctness of this sounding appears doubtful, however. In April, 1941, the Russian aviator Cherevichny, who landed on the ice in three different localities to the north of Wrangel Island and spent from three and a half to six days in each place, obtained much smaller depths (unpublished data communicated through the American Russian Institute, San Francisco, California). Cherevichny's soundings are as follows:

Latitude, north Longitude Depth, m
81°02ʹ 180°00ʹE 2647
78 30 176 40 E 1856
80 00 170 00 W 3430

These soundings were not available when the bathymetric chart of the Arctic region (fig. 3) was prepared.

The more or less elliptical North Polar Basin is connected with the Norwegian Basin by a fairly deep channel between Greenland and Spitsbergen, in which the sill depth is about 1500 m (table 6). The Norwegian Basin, in which there are two depressions with depths exceeding 3000 m, is separated from the open Atlantic by a ridge extending from Greenland to Scotland, from which Iceland and the Faeroe Islands rise above sea level. The sill depths in Denmark Strait, between Greenland and Iceland, and over the Wyville Thomson Ridge, between the Faeroes and Scotland, are about 500 m. Rising from the floor of the Norwegian Basin is an isolated elevation that extends above the surface as Jan Mayen Island.

Another depression of considerable magnitude that does not appear on chart I is in Baffin Basin between Baffin Island and Greenland, where the depths exceed 2000 m. This basin is separated from the open Atlantic by a ridge in Davis Strait between Baffin Island and Greenland, where the sill depth is about 700 m.


Interesting topographic features which are well developed in the Arctic regions are the “troughs” that cut across the shelf. These U-shaped furrows were apparently cut by glaciers at a period when the sea surface stood at a lower level. One such trough extends around the southern tip of Norway, and others may be traced by the irregularities of the 200-m contour to the north of Russia and between the islands of the Canadian Archipelago. Nansen (1928), in a discussion of the topography of the North Polar Basin, has described these features.


Polar projection showing the generalized sea-bottom topography of the Antarctic regions. Depths less than 4000 m shaded. Heavy dotted lines show the location of the elevations which separate the various basins. Contours at depths of 3000 m and more Correspond to those in chart I.

Fig. 4 is a polar projection of the Antarctic regions which shows the relationships between the major features of the submarine topography that cannot be visualized from chart I. The topography is based on the same data used for the preparation of chart I supplemented from other sources. All major depressions are also shown in the world map, but it has been possible to enter in this figure the contours above 3000 m. There are many striking differences between the topography of the North

and the South Polar regions. The semi-isolated basins that are so prominent in the North Polar region are lacking in the South. The shelf, with depths less than 200 m, is extensive only off the east coast of the southern part of South America, where the Falkland Islands rise from the shelf.

The deep basins extend close to the continent of Antarctica, and the slopes are relatively steep. Joining South America to Antarctica is the South Antilles Arc, upon which are located South Georgia, South Sandwich, South Orkney, and South Shetland Islands. The ridge is continuous at 4000 m, and at 3000 m there are only relatively narrow openings. The Atlantic Ridge does not extend as far south as Antarctica, but is terminated in the vicinity of Bouvet Island. The ridge to the south of Africa and Madagascar is known as the Crozet Ridge, after the island of that name that rises from it. Forming a part of the Indian Ridge is the conspicuous elevation surrounding Kerguelen Island known as the Kerguelen Ridge. The ridge extending from Australia to Antarctica supports Macquarie Island and is known as the Macquarie Ridge. The importance of these ridges in determining the distribution of properties and the character of the circulation around Antarctica is discussed in chapter XV, p. 610 et seq. The greatest depths found in the region shown in fig. 4 are in the Byrd Deep to the south of New Zealand and in the South Sandwich Trench on the convex side of the South Antilles Arc.

Bottom Configuration of Adjacent Seas

It is beyond the scope of this volume to present charts or descriptions of the many marginal and adjacent seas. In table 4 are listed the area, volume, and mean depth of some of these features. The adjacent seas of the Arctic regions are shown in fig. 3, and in figs. 5 and 6 are shown the generalized topographies of the European and American Mediterraneans. Details of the topography of other marginal areas are presented elsewhere. The degree of isolation—namely, the extent to which free exchange of water with the adjacent ocean is impeded by the presence of land or submarine barriers—plays an important role in determining the characteristic distribution of properties in such regions (see chapters IV and XV).

The European Mediterranean, which comprises the Mediterranean Sea, the Black Sea, and the waters connecting them (namely, the Dardanelles, the Sea of Marmora, and the Bosporus) forms an intercontinental sea bordered by Europe, Asia, and Africa. The Mediterranean Sea occupies a deep, elongated, irregular depression with an east-west trend, and the Black Sea occupies a small and topographically simpler depression offset to the north. The Black Sea Basin, with depths exceeding 2200 m, is virtually isolated from the Mediterranean Sea proper, the connection

being restricted to a narrow and shallow channel in the Bosporus where the sill depth is only 40 m, and a similar channel in the Dardanelles where the sill depth is 70 m. In the Sea of Marmora, depths exceed 1000 m. The Mediterranean Basin, where depths reach 4600 m, is in restricted communication with the open Atlantic through the Strait of Gibraltar, which is only 20 km wide and where the sill depth is about 320 m. The peculiar oceanographic conditions that prevail in the Black Sea and Mediterranean Sea (p. 642 et seq.) can be attributed to the isolation of the deeper waters.


Generalized bottom topography of the European Mediterranean. The larger basins are (I) Algiers-Provençal, (II) Tyrrhenian, (III) Ionian, (IV) Levantine, and (V) Black Sea Basin.

The generalized topography of the European Mediterranean is shown in fig. 5, which is based on a chart prepared by Stocks (1938). The Black Sea Basin (V) is of more or less elliptical form except in the north, where there are irregular shallow seas of which the largest is the Sea of Azov, east of the Crimean Peninsula. The connection with the Mediterranean Sea is through the Bosporus, the Sea of Marmora, and the Dardanelles into the Aegean Sea, where the irregular topography is reflected in the large number of islands. The Mediterranean Sea Basin is subdivided by a series of transverse ridges with a north-south trend, parts of which extend above sea level. The primary division into the western and eastern depressions is effected by a ridge extending from Europe to Africa—namely, Italy, Sicily, and the submerged part of the elevation between these land areas and Africa. The sill depth in the strait between Sicily and Tunis is about 400 m. The Western Mediterranean, in turn, is subdivided into the Algiers-Provençal Basin (I) and the Tyrrhenian Basin (II) by the ridge extending from northwestern Italy to Tunis, from which Corsica and Sardinia rise above the sea surface. The Eastern Mediterranean is subdivided into two major depressions: the Ionian Basin (III), in which maximum depths of 4600 m are found,

and the Levantine Basin (IV) by the ridge extending from Greece to Africa. There are other isolated depressions of smaller dimensions, such as the Alboran Basin between Spain and Morocco and those between Italy and Albania in the Adriatic Sea, to the north of Crete and to the south of Cyprus.


Generalized bottom topography of the American Mediterranean. The larger basins are (I) Mexico Basin; (II) Cayman Basin and (III) Cayman Trough, in the Western Caribbean; (IV) Colombia Basin and (V) Venezuela Basin, in the Eastern Caribbean. The greatest known depth in the Atlantic Ocean, 8750 m, is located in the Puerto Rico Trough to the north of Puerto Rico.

The American Mediterranean encompasses the partially isolated basins of the wide gulf bordered by North, Central, and South America which are separated from the open Atlantic by ridges, parts of which rise above sea level. The generalized topography of the region is shown in fig. 6, which is based on a chart by Stocks (1938). The chief difference between the European Mediterranean and the American Mediterranean is that the latter has numerous shallow and several deep connections with the open Atlantic. The topography of the American Mediterranean is extremely rugged, with deep trenches adjacent to steep-sided ridges, many of which rise above sea level. This is particularly true in the central and southern parts of the region, which are areas of pronounced gravity anomaly, volcanism, and strong seismic disturbances (Field et al, 1938). Bordering the low-lying coast of the Gulf of Mexico, off part of Honduras and Nicaragua, and surrounding the Bahama Islands are extensive shelves. The slopes leading down to deep water are in general rather steep, particularly between Cuba and Jamaica and along the

outer margin of the Bahama Islands. On the northern and outer side of the Lesser Antilles Arc is the Puerto Rico Trough, the deepest depression in the floor of the Atlantic Ocean, where the U.S.S. Milwaukee obtained a maximum sounding of 8750 m. Although there are many openings between the islands on the Lesser Antilles Arc, the openings with relatively great sill depths are few in number. In the Straits of Florida the sill depth is about 800 m, and in Windward Passage, between Cuba and Hispaniola, it is about 1600 m. A third deep channel with a sill at about 1600 m is located in Anegada and Jungfern Passages between the Virgin Islands and the Windward Islands and between the Virgin Islands and St. Croix Island. The oceanographic conditions (p. 640) indicate the presence of a small isolated basin between the Anegada and Jungfern Passages. With the exception of a channel between Dominica and Martinique with a saddle depth less than 1500 m, the Lesser Antilles Arc is continuous at depths less than 1000 m from the Windward Islands to the coast of South America.

The American Mediterranean is subdivided into two major depressions, the Gulf of Mexico and the Caribbean Sea, by a ridge between Yucatan and Cuba, and by the island of Cuba. The sill depth in the Yucatan Channel is less than 1600 m. The Mexico Basin (I) is a relatively simple depression lacking the irregularities that characterize the topography of the Caribbean region. Maximum depths of nearly 4000 m are found in the western part of the basin. The Caribbean region is separated into two major basins, the Western and the Eastern Caribbean, by the Jamaica Rise, which extends from Honduras to Hispaniola and from which Jamaica rises above the surface. The Western Caribbean is in turn divided into Yucatan Basin (II) and the Cayman Trough (III) by the Cayman Ridge, which extends westward from the southern extremity of Cuba. The Cayman Trough is the deepest depression in the American Mediterranean, and within the Bartlett Deep to the south of Cuba the U.S.S. S-21 obtained a maximum sounding of 7200 m. The Windward Passage between Cuba and Hispaniola appears to be a continuation of the depression forming the Cayman Trough. The greatest saddle depth between the Western and Eastern Carribbean is located in the passage between Jamaica and Hispaniola, where it is about 1200 m. The Eastern Caribbean is partially divided into two basins by the Beata Ridge, which extends south and west from Hispaniola toward South America. The western portion of the depression is known as the Colombia Basin (IV) and the eastern as the Venezuela Basin (V). The Aves Swell separates a small basin with depths greater than 3000 m in the eastern part of the Venezuela Basin, which is called the Grenada Trough.

The terminology to be applied to the features of the American Mediterranean is discussed by Vaughan in the report by Vaughan et al (1940), which also includes an excellent bathymetric chart of the Caribbean region prepared by the U. S. Hydrographic Office. The currents and distribution of properties in this area are described in chapter XV, p. 637.

Basin Max. depth (m) Adjacent deep depression Surface feature Location of sill Sill depth (m) Max. depth — sill depth (m)
Arctic Mediterranean Region
North Polar Basin 5400 North Pacific Bering Strait Siberia-Alaska 55 ....
Norwegian Basin Greenland-Spitsbergen 1500 3900
Norwegian Basin 3700 North Atlantic Denmark Strait Greenland-Iceland 500 3200
North Atlantic Faeroe Is.-Scotland 500 3200
Baffin Basin 2200 North Atlantic Davis Strait Baffin Is.-Greenland 700 1500
European Mediterranean Region
Western Mediterranean Basin 3700 North Atlantic Strait of Gibraltar Gibraltar-Morocco 320 3400
Eastern Mediterranean Basin 4600 Western Mediter-ranean Sicily-Tunis 400 4200
Black Sea Basin 2200 Eastern Mediter-ranean Bosporus 40
Dardanelles 70 2200
American Mediterranean Region
Eastern Caribbean Basin 5500 North Atlantic Anegada and Jungfern Passages Virgin Is.-Lesser Antilles 1600 3900
Western Caribbean Basin 7200 North Atlantic Windward Passage Cuba-Hispaniola 1600 5600
Eastern Caribbean Jamaica Channel Jamaica-Hispaniola 1200 ....
Mexico Basin 3900 Western Caribbean Yucatan Channel Yucatan-Cuba 1600 2300
North Atlantic Strait of Florida Florida-Bahama Is 800 ....
Other Regions
Japan Sea Basin 3700 Philippines Basin Tsushima Strait Korea-Japan 150 3550
Red Sea Basin 2800 Indian Ocean Strait of Bab-el-Mandeb Somaliland-Arabia 100 2700
Baltic Sea Basin 300 North Atlantic Danish Sounds Danish Is.-Germany 20 280


Isolated basins are of great interest from an oceanographic point of view, and in table 6 are brought together some of the data relating to the larger basins in adjacent seas. This tabulation does not include the basins in the East Indian Archipelago, which are discussed elsewhere (table 87, p. 738). The maximum depth within each basin and the greatest sill depth at which there is horizontal communication with the adjacent basins are listed as well as the difference between the greatest depth in the basin and the sill depth. The latter value corresponds to the depth of the “lake” that would be formed if the water level were lowered to the greatest sill depth. It will be seen that most of the basins listed are without horizontal communication through vertical distances of 3000 to 4000 m, and that in the Yucatan Basin the greatest depth is 5600 m below the sill. In great contrast to these deep basins is the Baltic Sea (average depth, 55 m), where depths greater than 300 m are restricted to small, isolated depressions and where the sill depth is only 20 m.

In fig. 7 is shown the topography of the area off the coast of Southern California. This coastal area is one of considerable interest in that it is physiographically similar to the adjacent land area and apparently represents a down-warped portion of the continent. The continental shelf is relatively narrow, and offshore is a series of basins and ridges upon which several islands are located. In the southern part the real continental slope leading down to the oceanic abyss is approximately 150 miles from the coast. This is not shown in the map. Several small canyons are also depicted in the figure.

Submarine Canyons

For many years it has been known that there were furrows cutting across the shelf in certain regions, but only since it became possible to obtain large numbers of accurately located soundings of the shelf and slope were such features found to be numerous and widespread. Variously termed canyons, valleys, mock-valleys, and gullies by different authors, they have stimulated a great deal of interest among geologists, and a large literature has been built up dealing with the character and mode of formation of these canyons. The data concerning the topography of the canyons have largely been obtained by national agencies engaged in the careful mapping of nearshore areas. Such data have been used by Veatch and Smith (1939) and by Shepard and Emery (1941) to prepare general and detailed topographic charts of the canyons off the east and west coasts of the United States. Stetson (1936) has carried out independent observations on the east coast, and Shepard

and his associates have made intensive studies of canyons, particularly on the west coast. The chief interest of the geologist has been to establish the manner in which the canyons were formed, and numerous hypotheses concerning their origin have been proposed. Shepard (Shepard and Emery, 1941) has discussed in detail the various hypotheses that have been advanced and offers arguments both for and against each of them.


Topography of a part of the very irregular sea bottom off the coast of southern California where there is a basin and ridge topography very similar to that on the adjacent land. Depths in fathoms (after Revelle and Shepard, 1939).

Although the terms listed above have been used more or less synonymously, the size and general character of the canyons vary greatly. Some of them off the mouths of rivers, such as the Hudson (fig. 9), Congo, and Indus Canyons, have depressions that can be traced across the shelf and even into the mouths of the rivers. Some canyons extend across the shelf, but others—for example, many of those shown in the charts prepared by Veatch and Smith—are limited to gashes in the continental slope and do not cut far across the shelf. The upper parts of the canyons are generally found to be steep-walled, V-shaped in profile, with the bottom sloping continuously outward (fig. 8). Some are winding, and many show a dendritic pattern, having smaller tributary canyons. In size they vary from small gullies to vast structures of the same dimensions as the Grand Canyon of the Colorado River (fig. 8).

The head of a canyon may sometimes be traced into shallow water within a few hundred meters of the land, and in other cases it is restricted to the upper part of the slope in depths of 100 to 200 m and at distances of fifty or more km from the coast. Certain canyons can be traced outward for great distances and into depths of several thousand meters, but Shepard is doubtful whether the deeper parts of the canyons, which are wider and have much gentler wall slopes, are of the same origin as the inner parts. The canyons are characteristic of continental coasts, but there is some evidence to show that similar features occur around oceanic islands.


Profiles of submarine canyons. (A) Transverse profile of the submarine canyon in Monterey Bay compared to a profile of the Grand Canyon of the Colorado River in Arizona (cf. fig. 10). (B) Transverse profiles of a small, steep-walled canyon off the southern California coast. (C) Longitudinal profiles of the Lydonia Canyon and the adjacent shelf and slope. (D) and (E) Transverse and longitudinal profiles of the Hudson Canyon, showing the relation to the adjacent shelf and slope. The locations of the transverse sections (D) are shown on the longitudinal profile. Note the vertical exaggeration in certain of the diagrams and the differences in horizontal scales (A and B after Shepard, 1938; C, D, and E after Veatch and Smith, 1939).

The steep walls of the canyons are generally free of unconsolidated sediment, and in those canyons where special investigations have been made the walls appear to be generally of sedimentary rock; in a few cases (for example, Monterey Canyon off the California coast, fig. 10) the canyons are cut into granite that is overlain by sedimentary rock. The sediments in the bottom of the canyons are generally coarser than those on the adjacent shelves, and in some of them cobbles and gravel have been found.


The following agencies have been advanced as possible causes for the formation of the canyons:

  1. Diastrophism.

  2. Erosion by submarine currents. Daly (1936) advanced the theory that “density currents” produced by suspension of much fine-grained sediment may have flowed down the slope and cut the canyons, particularly during intervals of lowered sea level during the glacial periods. Density currents occur in reservoirs, but there is no evidence of their existence in the sea, where the density stratification of the water impedes vertical flow.

  3. 42
  4. Spring sapping. Johnson (1939), in a thorough review of the literature concerning the character and origin of submarine canyons, develops the hypothesis that solution and erosion resulting from the outflow of underground water might contribute to the formation of the canyons.

  5. Mudflows and landslides. Mudflows are known to occur in the canyons (Shepard and Emery, 1941) and are agents which tend to keep the canyons clear of unconsolidated debris, but it is doubtful whether they are capable of eroding rock.

  6. Tsunamis or earthquake waves (p. 544). Bucher (1946) pointed out that most of the currents that might be found in canyons are of relatively low velocity and are therefore incapable of active erosion of rock. As a possible explanation of the submarine origin of the canyons he suggested that the rapid currents associated with earthquake waves set up in the sea by violent seismic motion of the sea bottom might be effective agents.

  7. Subaerial erosion. The five explanations listed above are compatible with the formation of the canyons below the sea surface. Because of their many resemblances to river-cut canyons on land, many investigators, notably Shepard, believe that the canyons must have had a subaerial origin. However, there is no accepted geological theory that would account for the world-wide exposure of the shelf and slope within relatively recent geological time. To overcome this difficulty, Shepard has suggested that during the ice ages the amount of water removed from the ocean and deposited as ice caps may have been much greater than ordinarily believed (p. 25).


Topography of the shelf and slope off part of the eastern coast of the United States showing different types of submarine canyons. The Hudson Canyon can be traced far across the shelf; others, such as the Lydonia, Oceanographer, and Hydrographer Canyons, cut into the outer margin of the shelf, while others are restricted to the slope itself. Depth contours in fathoms. (Simplified from chart in Veatch and Smith, 1939.)


Monterey Canyon off the coast of California.

Shepard (Shepard and Emery, 1941) has carefully evaluated the arguments in favor of and opposed to these various hypotheses concerning the origin of submarine canyons, and he concludes that no single hypothesis yet advanced can account for their characteristic features. Problems also exist concerning the processes which remove the sedimentary debris that must be swept into the canyons from the shelf. Mudflows and transportation by currents are known to be operative, but their effectiveness has not yet been determined.


The study of the development of shorelines has been carried out by geologists and physiographers who have classified the different types of coasts largely upon the basis of the extent to which erosion and deposition have affected the coastal configuration. Johnson (1919, 1925) has described the characteristic features of the coast and shallow-water zone, and these and other sources should be consulted in order to appreciate the complex nature of the transition zone between land and sea, where the effects of erosion and deposition, both subaerial and marine, must be

taken into account. In the chapter on marine sedimentation (chapter XX) the properties of beach and shallow-water sediments are discussed with reference to the source of sedimentary debris, the transporting agencies, and the roles of waves and currents in erosion and deposition. Although many observational data have been obtained on the features of the coastline, the evaluation of the relative importance of the different factors involved has been handicapped by the lack of adequate knowledge of the characteristics of waves, tidal currents, and other motions in the sea as transporting and eroding agents.

The classification of shorelines has been treated by Johnson (1919) and Shepard (1937). Shepard's primary classification is based upon a consideration of the following factors:

  1. Primary or youthful coasts with configurations due mainly to nonmarine agencies.

    1. Those shaped by terrestrial erosion agencies and drowned by deglaciation or down-warping.

    2. Those shaped by terrestrial depositional agencies such as rivers, glaciers, and wind.

    3. Those shaped by volcanic explosions or lava flows.

    4. Those shaped by diastrophic activity.

  2. Secondary or mature coasts with configurations primarily the result of marine agencies.

    1. Those shaped by marine erosion.

    2. Those shaped by marine deposition.

The beach is defined as the zone extending from the upper and landward limit of effective wave action to low-tide level. Consequently, the beach represents the real transition zone between land and sea, since it is covered and exposed intermittently by the waves and tides. The characteristics of beaches depend so much upon the nature of the source material composing their sediments and the effects of the erosion, transportation, and deposition by waves and currents that they can be more profitably discussed in the chapter on marine sedimentation. The upper part of the beach is covered only during periods of high waves, particularly when storms coincide with high spring tides. The slope of the beach is largely determined by the texture of the sediments (p. 1018), but the extent of the beach will depend upon the range in tide. The terminology applied to the various parts of the beach and the adjacent regions is shown in fig. 11, taken from a report by the Beach Erosion Board (U. S. Beach Erosion Board, 1933).

Beaches composed of unconsolidated material are characteristically regions of instability. Every wave disturbs more or less of the smaller sedimentary particles, and the character of the waves will determine whether or not there is a net removal or accretion of sediment in any

locality during a given time interval. Currents that effect a net transport of material kept in suspension by the waves also play an important part in shaping the beach. Fluctuations in the direction and height of the waves or in the direction of the currents alongshore usually result in changes in the profile of the beach. Such changes are commonly seasonal, and corresponding changes occur in the amount of sand in any locality and in the profile of the beach (Shepard and LaFond, 1940). The instability of the beach over relatively short time intervals has many implications in connection with the beach as an environment for sedentary organisms (chapter VIII).


Terminology applied to various parts of the beach profile. Berms are small impermanent terraces which are formed by deposition during calm weather and by erosion during storms. The plunge point is the variable zone where the waves break, hence its location depends on the height of the waves and the stage of the tide.

Although subject to short-period disturbances, the beach in general represents an equilibrium condition, despite the slow erosion of the coast or the permanent deposition that may be taking place. If the normal interplay of waves and currents is impeded in any way, as by the building of piers, breakwaters, or jetties, the character of the beach may be entirely changed. In some instances, highly undesirable erosion of the coast may result, and in others equally undesirable deposition may result. These changes will proceed until a new equilibrium is established that may render the value of the structure worthless for the purpose for which it was originally intended. The construction of breakwaters, jetties, sea walls or groins, and similar structures on an open coast should be undertaken only after a careful investigation of the character and source of the sedimentary material, the prevailing currents, the strength and direction of the waves, and other factors that determine the equilibrium form of the beach. The Beach Erosion Board of the U. S. Army, Corps of Engineers, as well as various private organizations, are engaged in studies of this type.



Bencker, H.1930. “The bathymetric soundings of the oceans” . Hydrographic Review, v. 7, no. 2, p. 64–97, 1930. Monaco.

Bucher, Walter H.1933. The deformation of the earth's crust. Princeton Univ. Press, 518 pp., 1933.

Bucher, Walter H.1940. “Submarine valleys and related geologic problems of the North Atlantic” . Geol. Soc. Amer., Bull., v. 51, p. 489–512, 1940.

Daly, Reginald A.1934. The changing world of the ice age. New Haven, Yale Univ. Press, 271 pp., 1934.

Daly, Reginald A.1936. “Origin of submarine “canyons.”” Amer. Jour. Sci., v. 31, p. 401–420, 1936.

Field, Richard M., et al.1938. “Symposium on the geophysical exploration of the ocean bottom arranged by the American Geophysical Union” . Amer. Philos. Soc., Proc., v. 79, p. 1–144, 1938.

Fowle, Frederick E.1933. “Smithsonian physical tables” . Smithsonian Misc. Coll, v. 88, 682 pp., 1933.

Gutenberg, Beno, ed. 1939. “Internal constitution of the earth” . v. 7 of Physics of the Earth. (Nat. Research Council.) New York, McGraw-Hill, 413 pp., 1939.

International Hydrographic Bureau. 1937. “Limits of oceans and seas” . Internat. Hydrogr. Bur., Spec. Pub. no. 23, 2d ed., 25 pp., 1937.

Johnson, Douglas W.1919. Shore processes and shoreline development. New York, Wiley & Sons, 584 pp., 1919.

Johnson, Douglas W.1925. The New England-Acadian shoreline, New York, Wiley & sons, 608 pp., 1925.

Johnson, Douglas W.1939. The origin of submarine canyons. New York, Columbia Univ. Press, 126 pp., 1939.

Johnstone, James. 1928. An introduction to oceanography. Liverpool, University Press, 368 pp., 1928.

Kossinna, Erwin. 1921. Die Tiefen des Weltmeeres. Berlin Univ., Institut f. Meereskunde, Veröff., N. F., A. Geogr.-naturwiss. Reihe, Heft 9, 70 pp., 1921.

Kuenen, Ph. H., 1935. “Geological interpretation of the bathymetrical results. Snellius Exped. in the eastern part of the Netherlands East Indies 1929–1930” , v. 5, Geological Results, pt. 1, 123 pp. and charts, 1935. Utrecht.

Littlehales, G. W.1932. “The configuration of the oceanic basins” . p. 13–46 in Physics of the Earth, v. 5, Oceanography. Nat. Research Council, Bull. no. 85, 1932. Washington, D. C.

Nansen, Fridtjof. 1928. “The oceanographic problems of the still unknown Arctic regions” . p. 3–14 in: Problems of Polar Research. Amer. Geogr. Soc., Spec. Pub. no. 7, 479 pp., 1928. New York.

Niblack, A. P.1928. “Terminology of submarine relief” . Hydrographic Review, v. 5, no. 2, p. 1–23, 1928. Monaco.

Raisz, Erwin. 1938. General cartography. New York, McGraw-Hill. 370 pp., 1938.

Revelle, Roger, and F. P. Shepard. 1939. “Sediments off the California coast” . p. 245–282 in: Recent Marine Sediments, A symposium. Amer. Assn. Petrol. Geol.736 pp., 1939. Tulsa.

Schott, G.1926. Geographie des Atlantischen Ozeans. Hamburg, C. Boysen, 368 pp., 1926.

Schott, G.1935. Geographie des Indischen und Stillen Ozeans. Hamburg, C. Boysen, 413 pp., 1935.


Shepard, Francis P.1931. “Glacial troughs of the continental shelf” . Jour. Geol., v. 39, p. 345–360, 1931.

Shepard, Francis P.1933. “Geological misconceptions concerning the oceans” . Science, v. 78, p. 406–408, 1933.

Shepard, Francis P.1937. “Revised classification of marine shorelines” . Jour. Geol., v. 45, p. 602–624, 1937.

Shepard, Francis P.1938. “Submarine canyons off the California coast” . Calif. Jour. Mines and Geol., Report 24, State Mineralogist, p. 298–310, 1938.

Shepard, Francis P.1939. “Near-shore sediments—hemipelagic deposits” . p. 219–229 in: Recent Marine Sediments, A symposium. Amer. Assn. Petrol. Geol., 736 pp., 1939. Tulsa.

Shepard, Francis P.1941. Unpublished data.

Shepard, F. P., and K. O. Emery. 1941. “Submarine topography off the California coast: canyons and tectonic interpretations” . Geol. Soc. Amer., Spec. Paper, no. 31, 171 pp., 4 charts, 1941.

Shepard, F. P., and E. C. La Fond. 1940. “Sand movements along the Scripps Institution pier” . Amer. Jour. Sci., v. 238, p. 272–285, 1940.

Stetson, Henry C.1936. “Geology and paleontology of the Georges Bank canyons” . Geol. Soc. Amer., Bull., v. 47, p. 339–366, 1936.

Stocks, Theodor. 1938. “Morphologie des Atlantischen Ozeans. Statistik der Tiefenstufen des Atlantischen Ozeans” . Deutsche Atlantische Exped., Meteor, 1925–1927, Wiss. Erg., Bd. 3, 1. Teil, 2. Lief., p. 35–151, 1938.

Stocks, Theodor, and G. Wüst. 1935. “Die Tiefenverhältnisse des offenen Atlantischen Ozeans” . Deutsche Atlantischen Exped. Meteor, 1925–1927, Wiss. Erg., Bd. 3, Teil 1, l., Lief., 31 pp., 1935.

U. S. Beach Erosion Board. 1933. Interim report. April 15, 1933.

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Veatch, A. C., and P. A. Smith. 1939. “Atlantic submarine valleys of the United States and the Congo Submarine Valley” . Geol. Soc. Amer., Spec. Pap. no. 7, 101 pp., 1939.

Wüst, Georg. 1936. “Die Gliederung des Weltmeeres” . Hydrographic Review, v. 13, no. 2, p. 46–56, 1936. Monaco.


III. Physical Properties of Sea Water

The properties of pure water are unique in comparison with those of other liquids, and the nature of our physical environment—that is, the characteristics of the oceans, the atmosphere, and the land—is in many ways dependent upon the peculiar properties of water. The “fitness” of water for the manifold needs of living organisms has been pointed out by physiologists and ecologists (for example, Henderson, 1913, Bayliss, 1927). Table 7 shows some of the characteristics which are important in this respect for both fresh water and sea water.

The unique character of water is further emphasized by the fact that, according to studies on related compounds, pure water should freeze at about − 150°C and boil at − 100°C. The chemical constitution of water offers no explanation for these anomalies and certain other deviations from the theoretical properties. However, it has been found that liquid water is not made up of individual H2O molecules but that it is polymerized—that is, multiple groups containing one, two, or three elementary H2O molecules may exist. These are referred to as monohydrol, dihydrol, and trihydrol. The relative proportions of the three forms depend upon the temperature, the immediate past history of the water, and other factors. The degree of polymerization decreases with increasing temperature. The existence of the water polymers is helpful in explaining certain of the peculiar properties of water such as the high melting and boiling points. The degree of polymerization has been thought to have certain physiological effects (Barnes and Jahn, 1934).

The discovery of isotopes of hydrogen and oxygen has modified our conception of “pure” water. All naturally occurring water contains small but variable amounts of heavy hydrogen (deuterium) and heavy oxygen. These modify the density and other properties, but, since their concentrations are extremely small, the effect is slight. As the understanding of this problem increases, it may become desirable to define more exactly certain physical units that are based on the properties of water, such as the liter and calorie, by taking into account the isotopic composition of water. The amounts of the heavy isotopes vary somewhat, depending upon the source of the water. Wirth, Thompson, and Utterback (1935) found that distilled water prepared from sea-water samples was, on the average, 1.4 × 10−6 greater in density than distilled tap water. Sea water from areas of great dilution showed a somewhat smaller anomaly—namely, 0.2 to 0.5 × 10−6 in density. Samples from the depths of the ocean were higher than the average. Swartout and Dole (1939) found that the density of water distilled from a sea-water sample was 1.7 × 10−6 greater than that prepared from Lake Michigan water. They also found that the ratio between hydrogen and deuterium in sea water was about 7000:1. Rain water, and consequently fresh water, has a lower proportion of the heavy isotopes than sea water, since a certain fractionation takes place in the process of evaporation. The field of isotope study is relatively new, and as yet little is known about the distribution or significance of the isotopes. Their variability is apparently rather small and consequently will not significantly affect the magnitude of the properties to be discussed.

CERTAIN PHYSICAL PROPERTIES OF WATER (In part after Fleming and Revelle, 1939)
Property Comparison with other substances Importance in physical-biological environment
Heat capacity Highest of all solids and liquids except liquid NH3 Prevents extreme ranges in temperature
Heat transfer by water movements is very large
Tends to maintain uniform body temperatures
Latent heat of fusion Highest except NH3 Thermostatic effect at freezing point owing to absorption or release of latent heat
Latent heat of evaporation Highest of all substances Large latent heat of evaporation extremely important in heat and water transfer of atmosphere
Thermal expansion Temperature of maximum density decreases with increasing salinity. For pure water it is at 4°C Fresh water and dilute sea water have their maximum density at temperatures above the freezing point. This property plays an important part in controlling temperature distribution and vertical circulation in lakes
Surface tension Highest of all liquids Important in physiology of the cell
Controls certain surface phenomena and drop formation and behavior
Dissolving power In general dissolves more substances and in greater quantities than any other liquid Obvious implications in both physical and biological phenomena
Dielectric constant Pure water has the highest of all liquids Of utmost importance in behavior of inorganic dissolved substances because of resulting high dissociation
Electrolytic dissociation Very small A neutral substance, yet contains both H+ and OH ions
Transparency Relatively great Absorption of radiant energy is large in infra red and ultraviolet. In visible portion of energy spectrum there is relatively little selective absorption, hence is “colorless.” Characteristic absorption important in physical and biological phenomena
Conduction of heat Highest of all liquids Although important on small scale, as in living cells, the molecular processes are far outweighed by eddy conduction (see text)


Information as to the physical properties of distilled water is found in comprehensive handbooks (for example, Dorsey, 1940) or in the International Critical Tables. A number of these physical properties depend upon two variables, temperature and pressure, but in the case of sea water a third variable has to be considered; namely, the salinity of the water, which will be defined and discussed below. Some of the properties, such as compressibility, thermal expansion, and refractive index, are only slightly altered by the presence of dissolved salts, but other properties that are constant in the case of distilled water, such as freezing point and temperature of maximum density, are dependent on salinity in the case of sea water. Furthermore, the presence of dissolved salts adds a few new characteristics to sea water, such as osmotic pressure. General surveys of the physical properties of sea water have been given by Krümmel (1907), Matthews (1923), Harvey (1928), Johnstone (1928), and Thompson (1932).

Another important aspect of the physical characteristics that has to be considered when dealing with water as it occurs in nature, regardless of whether fresh water or sea water is concerned, is that several important processes are greatly modified by the presence of minute suspended particles or by the state of motion. Thus, the absorption of light in lakes or in the sea is entirely different from the absorption of light in distilled water or in “pure” sea water, because the waters encountered in nature always contain suspended matter that causes increased scattering of the light and consequently increased absorption in layers of similar thickness. The processes of heat conduction, chemical diffusion, and transfer of momentum from one layer to another are so completely altered in moving water that, for water under natural conditions, the coefficients which have been determined under laboratory conditions must be replaced by corresponding “eddy” coefficients that depend

upon the presence of eddies. Some of the physical properties of sea water, therefore, depend only upon the three variables, temperature, salinity, and pressure, which can all be determined with great accuracy, whereas others depend upon such variables as amount of suspended matter or character of motion, which at present cannot be accurately determined. Before turning to a discussion of the physical properties and their relations to these variables, we shall discuss the salinity of the sea water.

Salinity and Chlorinity

In the chapter on chemical oceanography the composition of the dissolved constituents in sea water is considered in detail. Our present interest is only in the more abundant substances whose concentration will affect the physical properties. It has been found that, regardless of the absolute concentration, the relative proportions of the different major constituents are virtually constant, except in regions of high dilution (low salinity), where minor deviations may occur. From this rule it follows that any one of the major constituents may be used as a measure of the others and of the total amount of salt, and that water samples having the same total salt content, regardless of their source, are virtually identical in their physical properties.

Owing to the complexity of sea water, it is impossible by direct chemical analysis to determine the total quantity of dissolved solids in a given sample. Furthermore, it is impossible to obtain reproducible results by evaporating sea water to dryness and weighing the residue, because certain of the materials present, chiefly chloride, are lost in the last stages of drying. These difficulties can be avoided by following a technique yielding reproducible results which, although they do not represent the total quantity of dissolved solids, do represent a quantity of slightly smaller numerical value that is closely related and by definition is called the salinity of the water. This technique was established by an International Commission (Forch, Knudsen, and Sörensen, 1902), and on the basis of its work the salinity is defined as the total amount of solid material in grams contained in one kilogram of sea water when all the carbonate has been converted to oxide, the bromine and iodine replaced by chlorine, and 'all organic matter completely oxidized.

The determination of salinity by the method of the International Commission is rarely if ever carried out at the present time because it is too difficult and slow, but, owing to the constant composition of the dissolved solids, the determination of any of the elements present in relatively large quantity can be used as a measure of the other elements and of the salinity. Chloride ions make up approximately 55 per cent of the dissolved solids and can be determined with ease and accuracy by titration with silver nitrate, using potassium chromate as indicator.

The empirical relationship between salinity and chlorinity, as established by the International Commission, is
The chlorinity that appears in this equation is also a defined quantity and does not represent the actual amount of chlorine in a sample of sea water. Both salinity and chlorinity are always expressed in grams per kilogram of sea water—that is, in parts per thousand, or per mille, for which the symbol ‰ is used.

Chlorinity. In the titration with silver nitrate, bromides and iodides are precipitated together with the chlorides, but in the computation it is assumed that they are chlorides. Chlorinity was therefore originally defined as the total amount of chlorine, bromine, and iodine in grams contained in one kilogram of sea water, assuming that the bromine and the iodine had been replaced by chlorine. This definition introduces a chlorine-equivalent that is dependent upon the atomic weights used in preparing the standard solutions. Since the time of the work of the International Commission, there have been changes in the atomic weights, and the relation between salinity and chlorinity as defined above is no longer strictly true. In order to retain this relationship and to avoid apparent changes in the chlorinity of sea water it has fortunately been possible to redefine chlorinity so that it is independent of changes in atomic weights.

The primary standard used in the determination of chlorinity is so-called “Normal Water” (Eau de mer normale), prepared by the Hydrographical Laboratories in Copenhagen, Denmark, and distributed to all oceanographic institutions. Some of these institutions made their own secondary standards by means of the Normal Water. As a result of world conditions the preparation of Normal Water has temporarily been taken over by the Woods Hole Oceanographic Institution.

Normal Water is sea water whose chlorinity has been adjusted to about 19.4 ‰ and accurately determined by either direct or indirect comparisons with the original standard prepared in 1902. Hence, the chlorinities of all batches have been independent of changes in the atomic weights. A new primary standard (Urnormal-1937), prepared in 1937, will be used to establish the chlorinity of future batches of Normal Water for general distribution (Jacobsen and Knudsen, 1940). Comparison with earlier series of Normal Water showed the chlorinity of the new standard to be 19.381 ‰. As a more absolute method should be available with which to check future Normal Water preparations, the 1937 primary standard was analyzed, using extremely pure “atomic weight silver.” The equivalent amount of silver necessary to precipitate the halides was determined and the ratio of chlorinity to silver was found to be 0.3285233. On this basis a new definition of chlorinity was

introduced: The number giving the chlorinity in grams per kilogram of a sea-water sample is identical with the number giving the mass in grams of “atomic weight silver” just necessary to precipitate the halogens in 0.3285233 kilogram of the sea-water sample. By this redefinition the chlorinity has been made independent of changes in atomic weights, whereas the chlorine-equivalent, for which the original definition is retained, may vary slightly as the atomic weights are modified. The empirically established relationship between chlorinity and salinity (p. 51) remains valid and, according to the new definition of chlorinity, will never be affected by modification of the atomic weights.

The ratio of chlorine to silver, using the 1940 atomic weights (Ag = 107.880 and Cl = 35.457), is 0.3286707. Hence the ratio of chlorine-equivalent to the chlorinity is:

This ratio is important when computing the chemical composition of sea water of given chlorinity, as the chlorine-equivalent will be greater than the indicated chlorinity. It must also be taken into account when preparing standard chloride solutions or when direct gravimetric analyses are made to determine the halide content of sea water. It is of interest to see how closely this ratio corresponds to the change which could be expected from the modification of the atomic weights. According to Jacobsen and Knudsen (1940) the indicated chlorinity of the Normal Water should be increased by 0.0094 ‰ to take this modification into account. This procedure gives a ratio between chlorine-equivalent and chlorinity of 1.000485, which is in good agreement with that given above.

When dealing with the chemistry of sea water, other substances are generally determined and reported on a volume basis, wherefore it is convenient to introduce chlorosity (Cl) (Intern. Assn. Phys. Oceanogr., 1939), which is the property corresponding to the chlorinity expressed as grams per 20°-liter (p. 169). Chlorosity is obtained by multiplying the chlorinity of a water sample by its density at 20°. Table 8 gives the corresponding values of chlorosity for chlorinities between 15.00 and 21.00 ‰.

Methods for Obtaining Salinity, Other than Titration with Silver Nitrate. The salinity can also be determined from the density of a water sample at a given temperature or by measuring either the electrical conductivity or the refractive index, both of which depend upon the salinity. The character of these properties will be dealt with below, but their application to salinity determinations will be briefly discussed here.

Determinations of density are appropriately included under methods for obtaining salinity, because such determinations are generally made at

atmospheric pressure and at room temperature and will thus differ from, the density of the water sample at the locality where it was collected, but from the density thus observed the density at 0°C is computed, and from the latter the salinity can be found by means of Knudsen's Hydrographical Tables (p. 56). Determinations of density are rarely made, as it is difficult and time consuming to obtain an accuracy comparable to that obtained from the chlorinity titration, which is about ±0.00001. The methods used can be classified in two groups. In one, the mass of an accurately known volume of water at a definite temperature is determined, for example, by using a pycnometer bottle. In the second group, some form of hydrometer or float is used, and the density is computed from the weight of the hydrometer and the volume of the displaced water. The common form of stem hydrometer is generally not sufficiently accurate except when dealing with coastal waters, where great differences in density are found in short distances. The chain hydrometer of Hans Pettersson (1929) gives greater accuracy, but not as high as desirable. Nansen (1900) developed a hydrometer of total immersion which is very sensitive but which requires a water sample of at least 300 ml. The buoyancy is adjusted by the addition or removal of small weights until the hydrometer neither sinks nor rises in the sample, the temperature of which must be known within ±0.02°. Another type of instrument makes use of a sinker suspended in the water sample from one arm of a delicate chemical balance (Cummings, 1932). The “weight” of the sinker in a sample at a known temperature is determined, and from this the density may be computed. In all instances various corrections must be applied and the original sources consulted before any such determinations are attempted. Thompson (1932) has described the methods in some detail and gives many references.

Chlorinity, ‰ 15.00 16.00 17.00 18.00 19.00 20.00 21.00
Chlorosity, g/L 15.28 16.32 17.37 18.41 19.46 20.51 21.57
.28 .32 .37 .41 .46 .51 .57

Measurements of the electrical conductivity have been employed on board the Carnegie and by the U. S. Coast Guard as the routine method for obtaining salinities. Owing to the relatively high concentration of the ions and the effect that temperature has upon the conductivity, the apparatus and the technique employed are rather complicated (Wenner, Smith, and Soule, 1930; Soule, 1932). These instruments are standardized empirically, using sea-water samples of known salinity (determined by silver nitrate titrations against Normal Water), and the values for the unknown samples are obtained by interpolation. In order to obtain

results of adequate accuracy, extreme care must be taken to control the temperature of the conductivity cells, and the resistance must be measured very accurately. A simplified diagram of the type of circuit used in electrical conductivity measurements on sea water is shown in fig. 12. A and B represent two similar electrolytic cells. R is a variable resistance in series with B, the cell used for the unknown sample. A contains sea water of known chlorinity or a standard potassium chloride solution of approximately the same conductance. C and D are fixed resistances and S is a slide-wire resistance. H is a source of alternating current of frequency between 600 to 1000 cycles per second at a potential of about 0.5 to 1.0 volt. T represents the telephone receiver used to establish the balance of the bridge. Two cells are used to eliminate small temperature effects. For details concerning the circuit and the various instruments the original references cited above should be consulted.

According to Thomas, Thompson, and Utterback (1934) the Grinnel Jones conductivity bridge may be used with Washburn pipette-type cells, the constants of which are determined with standard potassium chloride solutions.


Circuit used in measurements of the electrical conductivity as a means of obtaining the salinity of sea water. Symbols are explained in the text.

The refractive index of sea water, which will be discussed later (p. 70) in greater detail, varies only slightly within the ranges of temperature and salinity encountered in the sea, but, in an interferometer, differences in refractive index can be measured with extreme accuracy. At a given temperature such differences depend only upon the salinity, and special types of interferometers have therefore been developed for indirect determinations of salinity, using water of known salinity as a standard. Monochromatic light must be used because the refractive index varies with the wave length. This method of determining salinity has not been widely used.

Units of Temperature, Salinity, and Pressure, and Their Ranges in the Sea

In oceanography the temperature is measured in degrees centigrade. The thermometers used are described on p. 347. The accuracy of the measurements is about ± 0.02°C. Salinity is given as grams per kilogram of sea water; that is, in parts per thousand, or per mille, for which the symbol ‰ is used. An accuracy of ±0.02 ‰ is required. Pressure (p. 170) is measured in atmospheres or in units of the c.g.s. system. An atmosphere is defined as the pressure exerted per square centimeter by a

column of mercury 760 mm high at a temperature of 0°C, where the acceleration of gravity is 980.665 cm/sec2. In chemical oceanography a related unit, the Torr, is used which equals the pressure exerted per square centimeter by a column of mercury 1 mm high at a temperature of 0°C and at the above-mentioned acceleration of gravity. The c.g.s. unit of pressure is dyne/cm2, and 1 atmosphere = 1.0133 × 106 dynes/cm2. One million dynes/cm2 was designated as 1 bar by V. Bjerknes. The corresponding practical unit used in physical oceanography is 1 decibar, which equals 0.1 bar. The pressure exerted per square centimeter by 1 m of sea water very nearly equals 1 decibar; that is, the hydrostatic pressure in the sea increases by 1 decibar for approximately every meter of depth. Therefore, the depth in meters and the pressure in decibars are expressed by nearly the same numerical value. This rule is sufficiently accurate for determining the effect of pressure on the physical properties of the water, but details of the pressure distribution must be computed from the density distribution (p. 408).

In the oceans the temperature ranges from about −2° to +30°C. The lower limit is determined by the formation of ice, and the upper limit is determined by processes of radiation and exchange of heat with the atmosphere (p. 127). In landlocked areas the surface temperature may be higher, but in the open ocean it rarely exceeds 30°C.

The salinity in the oceans is generally between 33 ‰ and 37 ‰. The surface salinity in high latitudes, in regions of high rainfall, or where there is dilution by rivers may be considerably less, and in certain semi-enclosed areas, such as the Gulf of Bothnia, the salinity may approach zero. In isolated seas in intermediate latitudes, such as the Red Sea, where evaporation is excessive, salinities may reach 40 ‰ or more. As the range in the open oceans is rather small, it is sometimes convenient to use a salinity of 35 ‰ as an average for all oceans. In the chapter on the chemistry of sea water the tabulations are made for water of 19.00 ‰ chlorinity; that is, 34.325 ‰ salinity.

In dealing with the pressure in the oceans, the atmospheric pressure is always neglected and the pressure at the sea surface is entered as zero. Since the pressure is essentially a function of depth and the numerical value in decibars nearly equals the depth in meters, the range in pressure will be from zero at the sea surface to over 10,000 decibars in the deepest part of the ocean.

Owing to the character of the distribution of temperature and salinity in the oceans, some relationships exist between these conditions and the pressure. The temperature of the deep and bottom water of the oceans is always low, varying between 4° and − 1°C, so that high pressures are associated with low temperatures. Similarly, the salinity of deep and bottom water varies within narrow limits, 34.6 ‰ to 35‰, and high pressures are therefore associated with salinities between these limits.

Exceptions are found in isolated seas in intermediate latitudes, such as the Mediterranean and Red Seas, where water of high temperature and high salinity is found at great depths, and hence under great pressure.

Density of Sea Water

The density of any substance is defined as the mass per unit volume. Thus, in the c.g.s. system, density is stated in grams per cubic centimeter. The specific gravity is defined as the ratio of the density to that of distilled water at a given temperature and under atmospheric pressure. In the c.g.s. system the density of distilled water at 4°C is equal to unity. In oceanography, specific gravities are now always referred to distilled water at 4°C and are therefore numerically identical with densities. In oceanography the term density is generally used, although, strictly speaking, specific gravity is always considered.

The density of sea water depends upon three variables: temperature, salinity, and pressure. These are indicated by designating the density by the symbol ρs,ϑ,p, but, when dealing with numerical values, space is saved by introducing σs,ϑ,p which is defined in the following manner:

Thus, if ρs,ϑ,p = 1.02575, σs,ϑ,p = 25.75.

The density of a sea-water sample at the temperature and pressure at which is was collected, ρs,ϑ,p is called the density in situ, and is generally expressed as σs,ϑ,p. At atmospheric pressure and temperature ϑ°C, the corresponding quantity is simply written σt, and at 0° it is written σ0. The symbol ϑ will be used for temperature except when writing σt, where, following common practice, t stands for temperature.

At atmospheric pressure and at temperature of 0°C the density is a function of the salinity only, or, as a simple relationship exists between salinity and chlorinity, the density can be considered a function of chlorinity. The International Commission, which determined the relation between salinity and chlorinity and developed the standard technique for determinations of chlorinity by titration, also determined the density of sea water at 0° with a high degree of accuracy, using pycnometers. From these determinations the following relation between (σ0 and chlorinity was derived:

Corresponding values of σ0, chlorinity, and salinity are given in Knudsen's Hydrographical Tables for each 0.01 ‰ Cl

In order to find the density of sea water at other temperatures and pressures, the effects of thermal expansion and compressibility on the density must be known. The coefficient of thermal expansion has been determined in the laboratory under atmospheric pressure, and

according to these determinations the density under atmospheric pressure and at temperature ϑ° can be written in the form
The quantity D is expressed as a complicated function of σ0 and temperature, and is tabulated in Knudsen's Hydrographical Tables. Since the values of σt are widely used in dynamical oceanography, tables for computing σt directly from temperature and salinity have been prepared by McEwen (1929) and Matthews (1932). A special slide rule for the same purpose has been devised by Sund (1929). Knudsen's tables also contain a tabulation of D as a function of σt and temperature, by means of which σ0 can be found if σt is known (σ0 = σt + D). This table is useful for obtaining the salinity of a water sample the density of which has been directly determined at some known temperature (p. 53).

The effect on the density of the compressibility of sea water of different salinities and at different temperatures and pressures was examined by Ekman (1908), who established a complicated empirical formula for the mean compressibility between pressures 0 and p decibars (quoted in V. Bjerknes and Sandström, 1910). From this formula, correction terms have been computed which, added to the value of σt, give the corresponding value σs,ϑ,p for any value of pressure.

Computation of Density and Specific Volume in Situ. Tables from which the density in situ, ρs,ϑ,p could be obtained directly from the temperature, salinity, and pressure with sufficiently close intervals in the three variables would fill many large volumes, but by means of various artifices convenient tables have been prepared. Following the procedure of Bjerknes and Sandström (1910), one can write

The first four terms can be expressed by σt, which can readily be determined by the methods outlined above, and the remaining terms represent the effects of the compressibility. When dealing with density it is desirable, for reasons that will be explained later (p. 402), to introduce the dynamic depth, D, as the independent variable instead of the pressure, p, and to write
The ∊ terms in this equation have been tabulated by Bjerknes and Sndström (1910) and by Hesselberg and Sverdrup (1914).

Instead of the density, ρs,ϑ,p, its reciprocal value, the specific volume in situ, αs,ϑ,p is generally used in dynamic oceanography. In order to avoid writing a large number of decimals, the specific volume is commonly expressed as an anomaly, δ, defined in the following way:


where α35,0,p is the specific volume of water of salinity 35 ‰, at 0°C, and at pressure p in decibars. The anomaly depends on the temperature, salinity, and pressure, and hence can be expressed as
It should be observed that the anomaly, by definition, does not contain a term δp, which would represent the effect of pressure at temperature 0° and salinity 35 ‰. The reason for this is explained on page 409. Of the above terms the last one, δs,ϑ,p is so small that it can always be neglected. Thus, five terms are needed for obtaining δ, and these were tabulated by Bjerknes and Sandström. If σt has already been computed, the terms that are independent of pressure can be combined as Δs,ϑ, (Sverdrup, 1933).

The value of Δs,ϑ, = δs + δϑ + δs, ϑ is easily obtained from σt because

Thus, in practice,
The values for these three terms are given in the appendix in small tables from which one can obtain the specific volume anomaly in situ of any water sample when its temperature, salinity, σt and the pressure are known. In these tables the terms are entered with one extra decimal place in order to avoid any accumulation of errors due to rounding-off of figures, and also in order to facilitate preparation of exact graphs that may be used instead of the tabulation, or for the preparation of tables in which the arguments are entered at such close intervals that interpolation becomes easy or unnecessary.

The procedure that is followed in calculating the density or specific volume in situ can be summarized as follows. For a given water sample the temperature, salinity, and depth at which it was collected must be known. For reasons stated elsewhere it can be assumed that the numerical value of the pressure in decibars is the same as that of the depth in meters. From the temperature and salinity the value σt is obtained from Knudsen's Tables or from graphs or tables prepared from this source, (McEwen, 1929; Matthews, 1932). With the values of σt temperature, salinity, and pressure the specific volume anomaly is computed by means of the tables given in the appendix. If the absolute value of the specific volume is required, the anomaly must be added to

the appropriate value of α35,0,p,ϑ, given in the appendix. In this table are given the specific volume of water of 35 ‰ and 0° at various pressures in decibars. The absolute density in situ can then be obtained as the reciprocal of the specific volume.

Another set of tables for computing the specific volume in situ has been prepared by Matthews (1938), who, in our notations, defines the anomaly as δ′ = αs,ϑ,p − α34,85,0,p. Thus, he refers the anomalies to water of salinity 34.85 ‰, for which σ0 = 28.00. The difference, δ − δ′ = α34,85,0,p − α35,0,p, depends upon the pressure:

Before comparing numerical values of the specific volume anomalies published in different reports, it is necessary to know on what tables the reported values are based.

Use of Knudsen's Hydrographical Tables. A certain point concerning the use of Knudsen's Hydrographical Tables should be kept in mind. Although they have been shown to hold very well over the normal range of the concentration of sea water, they are not necessarily valid for highly diluted or concentrated sea water. The tables are based on the careful examination of a series of samples collected from various regions. The dilute samples used were taken in the Baltic Sea, where dilution sometimes reduces the chlorinity to about 1 ‰, and where the river water that is mainly responsible for the dilution contains relatively large quantities of dissolved solids. This is shown by the fact that the equation relating salinity to chlorinity shows a salinity of 0.03 ‰ for zero chlorinity, and according to Lyman and Fleming (1940) the total dissolved solids corresponding to this figure are probably of the order of 0.07 ‰. Thus, empirically, the salinity of sea water can be expressed by an equation of the type

where the numerical value of a depends upon the composition of the diluted samples used for establishing the relation. If 1 kg of water of high salinity, S, is diluted by adding n kilograms of distilled water, the salinity of the dilution will be SD = S/(n + l), and the chlorinity of the diluted sample will be Cl/(n + 1). According to Knudsen's Tables this sample, however, has a salinity SK = a + b Cl/(n + 1). The difference between this and the true salinity is SK − SD = a[n/(n + l)], meaning that, if after dilution the chlorinity were determined by titration and the salinity were taken from Knudsen's Tables, it would be too high. Knudsen's Tables would therefore also give too great a density. As an example, let us assume that 1 kg of water of salinity 35 ‰ and chlorinity 19.375 ‰ is diluted by adding 9 kg of distilled water, reducing the
chlorinity to 1.938 ‰. Knudsen's Tables give for this chlorinity a salinity of 3.53 ‰, whereas the “true” salinity would be 3.50 ‰. Similarly, Knudsen's Tables would give a σ0 equal to 2.78, whereas the true value should be 2.75. At low concentration, chlorinities computed from direct density determinations, and vice versa, may therefore be in error. For example, “chlorinities” of sea ice computed from density measurements made on the melt water were consistently smaller than those determined by titration (p. 219), and in this case the diluting water was essentially distilled water. The restricted application of the Cl : S : density relations to highly diluted water occurring naturally or prepared in the laboratory should always be kept in mind.

Thermal Properties of Sea Water

Thermal Expansion. The coefficient of thermal expansion, e, defined by e = (l/αs,ϑ,p)(∂αs,ϑ,p/∂ϑ), is obtained, at atmospheric pressure, from the terms for D in Knudsen's Hydrographical Tables, and at higher pressures from Ekman's tables or formulae (p. 57). The coefficient for sea water is greater than that for pure water and increases with increasing pressure. A few numerical values are given in table 9, in which negative values indicate contraction with increasing temperature.

Pressure (decibars) Salinity ‰. Temperature (°C)
−2 0 5 10 15 20 25 30
0………… 0 −105 − 67 17 88 151 207 257 303
10 − 65 − 30 46 112 170 222 270 315
20 − 27 4 75 135 189 237 282 324
30 7 36 101 157 206 250 292 332
35 23 51 114 167 214 257 297 334
2,000………… 35 80 105 157 202 241 278
40 94 118 168 210 248 283
4,000………… 35 132 152 196 233 266
40 144 162 204 240 272
6,000………… 34.85 177 194 230
8,000………… 34.85 …… 231 246
10,000………… 34.85 …… 276 287

Thermal Conductivity. In water in which the temperature varies in space, heat is conducted from regions of higher to regions of lower

temperature. The amount of heat in gram calories per second which is conducted through a surface of area 1 cm2 is proportional to the change in temperature per centimeter along a line normal to that surface, and the coefficient of proportionality, κ, is called the coefficient of thermal conductivity (dQ/dt = −γ dϑ/dn). For pure water at 15°C the coefficient is equal to 1.39 ×10−3 The coefficient is somewhat smaller for sea water and increases with increasing temperature and pressure. This coefficient is valid, however, only if the water is at rest or in laminar motion (p. 89), but in the oceans the water is nearly always in a state of turbulent motion in which the processes of heat transfer are completely altered. In these circumstances the above coefficient of heat conductivity must be replaced by an “ceddy’ coefficient which is many times larger and which depends so much upon the state of motion that effects of temperature and pressure can be disregarded (p. 91).

Specific Heat. The specific heat is the number of calories required to increase the temperature of 1 g of a substance 1°C. When studying liquids, the specific heat at constant pressure, cp, is the property usually measured, but in certain problems the specific heat at constant volume, cv must be known.

The specific heat of sea water of different chlorinities was investigated by Thoulet and Chevallier, whose results have been recalculated and presented in different ways. Krümmel (1907) gives the following values for the specific heat at 17.5°C and atmospheric pressure:


Kuwahara (1939) gives an empirical equation for the specific heat at 0°C and atmospheric pressure:


It will be noted that the specific heat decreases with increasing salinity, but it has been pointed out by Krümmel that the effect is somewhat larger than might be expected from the composition of the solution, and the problem merits reinvestigation. The effects of temperature and pressure have not been measured, but it has been assumed that they are the same as those for pure water. Ekman (1914) gives the following values showing the effect of temperature on the specific heat of water of S = 34.85 ‰ at atmospheric pressure:


The effect of pressure on the specific heat has been computed by Ekman (1914) from the equation

where p is the pressure in decibars, T is the absolute temperature, ρ the density, J the mechanical equivalent of heat, and e the coefficient of thermal expansion. The combined effect of temperature and pressure is shown in table 10, where cp,0 is the specific heat at atmospheric pressure of water of salinity 34.85 ‰. The specific heat at constant volume, which is somewhat less than cp, may be computed from the following equation:
where K is the true compressibility (p. 68) and where the other symbols are defined above. According to Matthews (1923), the ratio of cp,:cv, for water of S = 34.85 ‰ at atmospheric pressure increases from 1.0004 at 0° to 1.0207 at 30°. The effect of pressure is appreciable; for the same water at 0°, the ratio is 1.0009 at 1000 decibars, and 1.0126 at 10,000 decibars. This ratio is important in the study of the velocity of sound (P. 76).

DIFFERENCE BETWEEN SPECIFIC HEAT AT ATMOSPHERIC PRESSURE AND AT PRESSURE p, (cp,0cp,p) (Salinity 34.85 ‰ at indicated temperature)
θ (°C) p (decibars)
2000 4000 6000 8000 10,000
− 2………… 0.0171 0.0315 0.0435
0………… 0.0159 0.0291 0.0401 0.0492 0.0566
5………… 0.0136 0.0248 0.0340 0.0416 0.0479
10………… 0.0120 0.0220
15………… 0.0110 0.0203
20………… 0.0105

Latent Heat of Evaporation. The latent heat of evaporation of pure water is defined as the amount of heat in gram calories needed for evaporating 1 g of water, or as the amount of heat needed for producing 1 g of water vapor of the same temperature as the water. Only in the latter form is the definition applicable to sea water. The latent heat of evaporation of sea water has not been examined, but it is generally assumed that the difference between that and pure water is insignificant; therefore, between temperatures of 0° and 30°C, the formula

can be used.


Adiabatic Temperature Changes. When a fluid is compressed, without gain or loss of heat to the surroundings, work is performed on the system and there is a rise in temperature. Conversely, if expansion takes place, the liquid itself gives up energy, which is reflected in a drop in temperature. Such adiabatic temperature changes are well known and important in the atmosphere. Sea water is compressible, and the effects of adiabatic processes, although small, must be taken into account when studying the vertical distribution of temperature in the great depths of the oceans and in deep isolated basins where the adiabatic heating may lead to a temperature increase toward the bottom (for example, p. 739). Adiabatic cooling is of immediate practical concern when water samples are taken with thermally insulated water bottles and the temperature of the water sample is determined after it has been raised to the surface (p. 355)

Depth (m) Temperature (°C)
−2 0 2 4 6 8 10 15 20
0. 0.016 0.035 0.053 0.078 0.087 0.103 0.118 0.155 0.190
1,000. 0.036 0.054 0.071 0.087 0.103 0.118 0.132 0.166 0.199
2,000. 0.056 0.073 0.089 0.104 0.118 0.132 0.146 0.177 0.207
3,000. 0.075 0.091 0.106 0.120 0.133 0.146 0.159 0.188
4,000. 0.093 0.108 0.122 0.135 0.147 0.159 0.170 0.197
5,000. 0.110 0.124 0.137 0.149
6,000. 0.120 0.140 0.152 0.163
7,000. …… 0.155 0.165 0.175
8,000. …… 0.169 0.178 0.187
9,000. …… 0.182 0.191 0.198
10,000. …… 0.194 0.202 0.209

Any adiabatic effect is related to changes in pressure, but in the sea the pressure can be considered proportional to the depth, and adiabatic temperature changes can be given as changes per unit of depth instead of per unit of pressure. According to Lord Kelvin, the change in temperature for each centimeter of vertical displacement is

where T is the absolute temperature and g is the acceleration of gravity, and where the other symbols have their previous meanings. This change
is extremely small, and for practical purposes the adiabatic temperature change on a vertical distance of 1000 m, called the adiabatic temperature gradient, is used instead. It should be observed that the adiabatic temperature gradient depends mainly upon the coefficient of thermal expansion, e, which varies much more with temperature and pressure than the other quantities involved. Ekman (1914) has computed the adiabatic temperature gradient for different temperatures, salinities, and depths, and some of his values are shown in table 11.

Depth (m) Temperature, ϑm (°C)
− 2 −1 0 1 2 3 4 5 6 7 8 9 10
1,000…… 2.6 3.5 4.4 5.3 6.2 7.0 7.8 8.6 9.5 10.2 11.0 11.7 12.4
2,000…… 7.2 8.9 10.7 12.4 14.1 15.7 17.2 18.8 20.4 21.9 23.3 24.8 26.2
3,000…… 13.6 16.1 18.7 21.2 23.6 25.9 28.2 30.5 32.7 34.9 37.1 39.2 41.2
4,000…… 21.7 25.0 28.4 31.6 34.7 37.7 40.6 43.5 46.3 49.1 51.9 54.6 57.2
5,000…… 31.5 35.5 39.6 43.4 47.2 50.9 54.4
6,000…… 42.8 47.5 52.2 56.7 61.1 65.3 69.4
7,000…… …… …… 66.2 71.3 76.2 80.9 85.5
8,000…… …… …… 81.5 87.1 92.5 97.7 102.7
9,000…… …… …… 98.1 104.1 109.9 115.6 121.0
10,000…… 115.7 122.1 128.3 134.4 140.2

The temperature that a water sample would attain if raised adiabatically to the sea surface has been called the potential temperature (Helland-Hansen, 1912b) and has been designated Θ. Thus, Θ = ϑm − Δϑ, where ϑm, is the temperature in situ and Δϑ is the amount by which the temperature would decrease adiabatically if the sample were raised to the surface. The potential temperature can be obtained from a table of the adiabatic gradients by step-wise computations. Such computations are long and tedious, but Helland-Hansen (1930) has prepared a, convenient table for Δϑ which has been reproduced in table 12. The table is based on a salinity of 34.85 ‰ (σ0 = 28.0) and is applicable in general to the deep ocean areas, because in these the salinity does not differ much from 34.85 ‰ and because the effect of salinity on the adiabatic processes is small. It may be seen that if water of 2°C is raised adiabatically from 8000 m to the surface, Δϑ = 0.925°, and therefore the potential temperature of that water is 1.075°. The adiabatic cooling of water of different

salinities that may occur near the surface and in the Mediterranean Sea is given in tables 13 and 14.

Colligative and Other Properties of Sea Water

Colligative Properties. The colligative properties—namely, vapor-pressure lowering, freezing-point depression, boiling-point elevation, and osmotic pressure—are unique properties of solutions. If the magnitude of any one of them is known for a solution under a given set of conditions, the others may readily be computed. Solutions of the complexity and concentration of sea water do not obey the generalized theories of the colligative properties, but in all cases the departures from the theoretical values are proportional.

Salinity (V%) Temperature, ϑm (°C)
0 2 4 6 8 10 12 14 16 18 20 22
30.0…… 3.5 5.3 7.0 8.7 10.3 11.8 13.2 14.7 16.1 17.6 18.9 20.3
32.0…… 3.9 5.7 7.3 9.0 10.6 12.1 13.5 15.0 16.4 17.8 19.1 20.5
34.0…… 4.3 6.0 7.7 9.4 10.9 12.4 13.8 15.3 16.6 18.0 19.3 20.7
36.0…… 4.7 6.4 8.1 9.7 11.2 12.7 14.1 15.5 16.9 18.3 19.6 20.9
38.0…… 5.1 6.8' 8.4 10.0 11.6 13.0 14.4 15.8 17.2 18.5 19.8 21.1
Depth (m) Temperature, ϑm (°C)
12 13 14
1000…………………… 14.4 15.1 15.8
2000…………………… 30.0 31.4 32.7
3000…………………… 46.6 48.6 50.6
4000…………………… 64.2 66.7 69.2

Only the depression of the freezing point for sea water of different chlorinities has been determined experimentally (Knudsen, 1903; Miyake,

1939a), and empirical equations for computing the vapor-pressure lowering and osmotic pressure have been based on these observations. Thompson (1932) has shown that the depressions of the freezing point, Δϑf, may be calculated from the chlorinity by means of the equation
The values of Δϑf for various chlorinities are shown in fig. 13. The freezing point of sea water is the “initial” freezing point; namely, the temperature at which an infinitely small amount of ice is in equilibrium with the solution. As soon as any ice has formed, the concentration of the dissolved solids increases, and hence the formation of additional ice can take place only at a lower temperature (p. 216).


Osmotic pressure, vapor pressure relative to that of pure water, freezing point, and temperature of maximum density as functions of chlorinity and salinity.

The vapor pressure of sea water of any chlorinity referred to distilled water at the same temperature can be computed from the following equation (Witting, 1908):

where e is the vapor pressure of the sample and e0 that of distilled water at the same temperature (fig. 13). Sea water within the normal range of concentration has a vapor pressure about 98 per cent of that of pure water at the same temperature, and in most cases it is not necessary to consider the effect of salinity, since variations in the temperature of the surface waters have a much greater effect upon the vapor pressure (table 29, p. 116).

The osmotic pressure can be calculated from the freezing-point depression by means of the equation derived by Stenius (Thompson, 1932):


The osmotic pressure at any temperature may then be computed:

The variations in osmotic pressure over the range in chlorinity from 5 ‰ to 22 ‰ are shown in fig. 13.

It will be noted that the freezing-point depression and, therefore, the other colligative properties are not linear functions of the chlorinity, because chlorinity is reported as grams per kilogram of sea water and not as grams per kilogram of solvent water, in which case a linear relationship should be expected. In agreement with this expectation, Lyman and Fleming (1940) found that the freezing-point depression could be written in the form

where Z is the total salt content per kilogram of solvent water.

The magnitude of the colligative properties depends upon the concentration of ions in the solution and upon their activity. According to present concepts the major constituents of sea water exist as ions whose concentrations may be computed from data in table 35 (p. 173). Within the normal range of sea water the gram-ionic concentration per kilogram of solvent water may be obtained from the following expression:

The gram-ionic concentration for water of 19 ‰ Cl is 1.1368. The gram-molecular lowering of the freezing point is −1.86°. Therefore, the “theoretical” value for the depression of the freezing point of water of chlorinity 19 ‰ should be −1.86 × 1.1368 = −2.11°, but the observed value for water of the same chlorinity is −1.872°. The ratio between the actual and the theoretical value, 0.89, is a measure of the reduced
activity of the ions in sea water of normal concentrations, and, as mentioned above, the other colligative properties bear the same relation to the theoretical values.

Maximum Density. Pure water has its maximum density at a temperature of very nearly 4°, but for sea water the temperature of maximum density decreases with increasing salinity, and at salinities greater than 24.70‰ is below the freezing point. At a salinity of 24.70 ‰, the temperature of maximum density coincides with the freezing point: ϑf = −1.332°. Consequently, the density of sea water of salinity greater than 24.70 ‰ increases continuously when such water is cooled to its freezing point. The temperature of maximum density is shown in fig. 13 as a function of salinity and chlorinity.

p (bars) Temperature (°C)
0 5 10 15 20 25 30
0 4659 4531 4427 4345 4281 4233 4197
100 4582 4458 4357 4278
200 4508 4388 4291
400 4368 4256
1000 4009 3916

Compressibility. Ekman (1908) has derived an empirical equation for the mean compressibility of sea water between pressures 0 and p bars (p. 57), as defined by αs,ϑ,p = αs,ϑ,0 (1 − kp). Numerical values are given in table 15, where the bar has been used as pressure unit.

The true compressibility of sea water is described by means of a coefficient that represents the proportional change in specific volume if the hydrostatic pressure is increased by one unit of pressure: K = (− l/α)(dα/dp). The true compressibility can be calculated from the mean compressibility, which was tabulated by Ekman, using the equation

where k is the mean compressibility referred to bar as pressure unit, and p is the pressure in bars.

Viscosity. When the velocity of moving water varies in space, frictional stresses are present. The frictional stress which is exerted on a

surface of area 1 cm2 is proportional to the change of velocity per centimeter along a line normal to that surface (τs, = μ dv/dn). The coefficient of proportionality (μ) is called the dynamic viscosity. This coefficient decreases rapidly with increasing temperature and increases slowly with increasing salinity (table 16, after Dorsey, 1940). With increasing pressure the coefficient for pure water decreases at low temperature but increases at high temperature (Dorsey, 1940). If the same holds true for sea water, and if the effect is of similar magnitude, the viscosity of water of salinity 35 ‰ and temperature 0° is 18.3 × 10−3 c.g.s. units at a pressure of 10,000 decibars, as against 18.9 × 10−3 at atmospheric pressure. The difference is insignificant, and the effect of pressure on the viscosity can be disregarded in the oceans.

Salinity (%) Temperature (°C)
0 5 10 15 20 25 30
0………… 17.9 15.2 13.1 11.4 10.1 8.9 8.0
10………… 18.2 15.5 13.4 11.7 10.3 9.1 8.2
20………… 18.5 15.8 13.6 11.9 10.5 9.3 8.4
30………… 18.8 16.0 13.8 12.1 10.7 9.5 8.6
35………… 18.9 16.1 13.9 12.2 10.9 9.6 8.7

The viscosity that has been discussed so far is valid only if the motion is laminar, but, as stated above, turbulent motion prevails in the sea, and an “eddy” coefficient must be introduced which is many times larger (p. 91).

Diffusion. In a solution in which the concentration of a dissolved substance varies in space, the amount of that substance which per second diffuses through a surface of area 1 cm2 is proportional to the change in concentration per centimeter along a line normal to that surface (dM/dt = −δ dc/dn). The coefficient of proportionality (δ) is called the coefficient of diffusion; for water it is equal to about 2 × 10−5, depending upon the character of the solute, and is nearly independent of temperature. Within the range of concentrations encountered in the sea the coefficient is also nearly independent of the salinity.

What was stated about the coefficient of thermal conductivity in the sea applies also to the coefficient of diffusion. Where turbulent motion prevails, it is necessary to introduce an “eddy” coefficient that is many times larger and that is mainly dependent on the state of motion.


Surface Tension. The surface tension of sea water is slightly greater than that of pure water at the same temperature. Krümmel (1907) carried out experimental observations from which he derived an empirical equation relating the surface tension to the temperature and salt content. This equation was revised by Fleming and Revelle (1939) to take into account the more recent values for pure water. The revised expression has the form

The surface tension is decreased by impurities, and in the sea is mostly smaller than stated.

Refractive Index. The refractive index increases with increasing salinity and decreasing temperature. The problem of determining the relationship between these variables, and the types of equipment to be used, has been discussed by a number of authors (for example, Utterback, Thompson, and Thomas, 1934; Bein, Hirsekorn, and Möller, 1935; Miyake, 1939). Since the index varies with the wave length of light, a standard must be selected, usually the D line of sodium.

Utterback, Thompson, and Thomas determined at a number of temperatures the refractive index of ocean-water samples that had been diluted with distilled water. They found that the refractive index could be represented by expressions of the following type:

where nϑ is the refractive index of the sea-water sample at the temperature ϑ°, n0,ϑ is that of distilled water at the same temperature, and kϑ is a constant appropriate for that temperature. This equation gives a straight-line relationship between the refractive index and the chlorinity, but it should be remembered that it is valid for ocean water diluted with distilled water and that at low chlorinities the diluted water does not correspond to sea water of the same low chlorinity, according to Knudsen's Hydrographical Tables. In fig. 14 the relationships between n, ϑ, and Cl determined by Utterback, Thompson, and Thomas are shown. Miyake (1939b) determined the refractive index for the sodium D line at 25° (nD,25°) for oceanic water samples that were diluted in the laboratory. He represented his results by the same type of equation, but obtained numerical constants that differ slightly from those of the authors mentioned above.
Miyake found that the refractive index of sea water could be expressed as

where n0 is the refractive index of distilled water tind v is the refractive index of solutions of single salts having concentrations comparable to those in which these salts occur in the sea water. It is known that individual ions have characteristic ionic refractions. In sea water the salts are completely ionized, and, as the molar refractions are known for each ion, Miyake was able to compute the refractive index with a fair degree of accuracy.

Electrical Conductivity. Thomas, Thompson, and Utterback (1934), and Bein, Hirsekorn, and Möller (1935) have studied conductivity as a function of chlorinity and temperature and have given tables for the specific conductance in reciprocal ohms per cubic centimeter for a wide range in conditions.


Refractive index of sea water as a function of temperature and chlorinity.

The results of the investigations of Thomas, Thompson, and Utterback are expressed at temperatures of 0, 5, 10, 15, 20, and 25°. Their results are shown graphically in fig. 15. The values for the low chlorinities were obtained by diluting ocean water with distilled water, and hence the density and other properties will not correspond exactly to those of water of the same chlorinity, as represented in Knudsen's Hydrographical Tables.

Properties of Sea Ice

The processes of freezing and the chemical properties of ice formed from sea water in high latitudes are discussed elsewhere. The physical properties of sea ice, like those of the water, depend upon the salt content, which in turn is a function of the rate of freezing, age, thermal history, and so forth. The salts in sea ice (p. 217) do not differ greatly in composition from those in the water, as they are generally present in brine that is enclosed in small cells. Therefore, within practical limits, the terms chlorinity and salinity of sea ice have the same meaning as for the water, although the salts are not uniformly distributed in the ice.

The freezing point of sea water, as was pointed out previously, represents the initial freezing point at which ice is in equilibrium with sea water of the indicated chlorinity. If the ice and sea water are in a closed system, as when brine is enclosed in cells in the ice, a further reduction of temperature is necessary to cause additional ice to separate.

Observations by Ringer (p. 217) make it possible to find the salt content of the brine that will be in equilibrium with ice at different temperatures (fig. 16A). The salt content used in this case is the total amount present, although at low temperatures certain salts crystallize out. The reason for giving the data in this way is that analyses for the chlorinity or salinity of the ice do not discriminate between that existing in solution or as crystals. Given the salt content of the brine that will be in equilibrium with the ice at any temperature, it is possible to calculate the amount of enclosed brine per kilogram of ice of unit salinity at any temperature. The amount of brine in any sample of ice at a given temperature can be obtained by multiplying the value at thet temperature, shown by the curve in fig. 16B, by the salinity of the ice. Thus, sea ice of salinity 10 ‰ at −3°C is essentially a mush containing 200 g of brine per kilogram. From the data in fig. 16 it is also possible to compute the amount of ice that is formed or melted when ice of a known salinity is cooled or heated.


Specific conductance, reciprocal ohms/cm3, of sea water as a function of temperature and chlorinity.

Another variable which markedly affects certain physical properties of sea ice is the gas content. The gases normally occur as small “bubbles” in the ice, and the quantity is generally large in ice that has frozen rapidly, in which case bubbles represent gases originally in solution in the water, or in old ice that has undergone partial thawing and been refrozen, in which case atmospheric air is trapped in the ice.


(A) Salt content of the enclosed brine in equilibrium with sea ice at different temperatures. (B) Amount of brine in 1 kg of sea ice of salinity 1 ‰ at different temperatures.

In the following discussion the numerical values relating to the properties of sea ice are quoted from the work of Malmgren (1927), unless otherwise noted. The corresponding values for pure ice represent physical constants for gas- and salt-free ice, and are taken from Barnes (1928).

Pure ice at 0° has a density of 0.91676, whereas pure water at the same temperature has a density of 0.9998674. The density of sea ice

varies both above and below that of pure ice, depending upon its content of water or brine and the gas content. Malmgren reports a range between 0.857, in old surface ice, and 0.92.

The specific heat of pure ice depends upon its temperature and varies within narrow limits, but that of sea ice is a much more variable property, depending upon the salt or brine content and the temperature. Changing the temperature of sea ice will generally involve either melting or freezing, and the amount of heat required will depend upon the salinity of the ice, as shown in table 17. It should be noted that the specific heat of pure ice is less than half that of pure water. Near the initial freezing point, the extremely high specific heat of ice of high salinity is, of course, due to the formation of ice from the enclosed brine or its melting.

Salinity ‰ Temperature (°C)
−2° −4° −6° −8° −10° −12° −14° −16° −18° −20° −22°
0………… 0.48 0.48 0.48 0.48 0.48 0.47 0.47 0.47 0.47 0.47 0.46
2………… 2.47 1.00 0.73 0.63 0.57 0.55 0.54 0.53 0.53 0.52 0.52
4………… 4.63 1.50 0.96 0.76 0.64 0.59 0.57 0.57 0.56 0.55 0.54
6………… 6.70 1.99 1.20 0.88 0.71 0.64 0.61 0.60 0.58 0.57 0.56
8………… 8.76 2.49 1.43 1.01 0.78 0.68 0.64 0.64 0.61 0.60 0.58
10………… 10.83 2.99 1.66 1.14 0.85 0.73 0.68 0.67 0.64 0.62 0.60
15………… 16.01 4.24 2.24 1.46 1.02 0.85 0.77 0.76 0.71 0.68 0.65

The latent heat of fusion of pure ice at 0°C and at atmospheric pressure is 79.67 calories per gram. As the melting of sea ice does not occur at a fixed temperature on account of the presence of the salts, it is impossible to designate the latent heat in the usual way. Malmgren gives the heat required to melt 1 g of ice of given salinity that was initially at the indicated temperature (table 18). It may readily be seen that the presence of salts reduces the apparent latent heat.

The vapor pressure of sea ice has not been investigated, but it cannot depart very much from that of pure ice, which has the following values:


The latent heat of evaporation of pure ice is variable. It has been found that under certain conditions the ice can volatilize directly to vapor without going through the liquid stage, in which case the latent heat of evaporation is about 600 calories per gram. If the evaporation

proceeds more slowly, the ice melts before vaporizing and 700 calories per gram are required. The latter process seems to prevail in nature.

Temperature (°C) Salinity (‰)
0 2 4 6 8 10 15
−1………… 80 72 63 55 46 37 16
−2………… 81 77 72 68 63 59 48

For pure ice the coefficient of thermal expansion (e) per one degree is about 1.7 × 10−4, where e = (l/α)(dα/dϑ). The coefficient is independent of temperature. The thermal expansion of sea ice is a function of its temperature and salinity and shows a considerable range over both positive and negative values, as shown in table 19, where negative values indicate expansion on cooling, positive values contraction on cooling. This anomalous behavior is again related to the salt or brine content because any change in temperature leads to the formation or melting of a certain amount of ice. Thus, the processes in sea ice are a combination of the sudden change in volume associated with the ice ⇌ water transformation and the thermal expansion of the brine and the ice. According to table 19, sea ice of high salinity expands rapidly as it is cooled below the initial freezing point. The coefficient decreases at lower temperatures but is always negative. On the other hand, ice of low salinity first expands and then contracts as its temperature is lowered.

The coefficient of thermal conductivity of pure ice is about 0.005, which is approximately three times as great as that of pure water at 0°. Malmgren carried out a number of measurements in the Arctic ice fields and found that the conductivity was greatly affected by the character of the ice, particularly by the gas content (that is, the porosity) of the ice. Under natural conditions, porosity is greater near the surface than in the deeper portions of the ice. On the average, the coefficient of thermal conductivity near the surface was about 1.5 × 10−3, at 0.5 m it was 4.0 × 10−3, and below 1.0 m it approached the value of pure ice given above—namely, 5.0 × 10−3.

Transmission of Sound

Water is a very efficient medium for the transmission of sound, which travels more rapidly and with much less absorption of energy through water than through air. This characteristic has made possible the development of submarine acoustic methods that are of tremendous value in navigation. The most familiar use is in echo sounding, where the time required for an impulse to travel to the sea floor and back to the vessel is used as a measure of the depth. Horizontal sound transmission is used in radio-acoustic range finding, which is employed by surveying vessels in order to determine their location accurately when carrying on sounding operations out of sight of land. The position of the vessel is determined by measuring the time intervals required for the impulse of a bomb explosion to travel to two or more submarine sound detectors (hydrophones) at known locations, usually near shore. Sound transmission from subsurface bells has also been used as a navigational aid near lighthouses. Ultrasonic impulses (frequency above the range audible to the human ear) are largely directional, and many attempts have been made to develop instruments for locating icebergs and other navigational hazards in the path of a vessel. The practical aspects of this problem and the types of equipment used are too numerous to consider in this discussion, and reference is made to the Hydrographic Review (Monaco) as an excellent source of material on these problems.

Salinity (‰) Temperature (°C)
−2 −4 −6 −8 −10 −12 −14 −16 −18 −20 −22
2………… −22.10 −4.12 −1.06 0.16 0.83 1.13 1.23 1.27 1.33 1.38 1.44
4………… −45.89 −9.92 −3.81 −1.37 −0.02 0.57 0.78 0.85 0.96 1.07 1.88
6………… −69.67 −15.73 −6.55 −2.90 −0.88 0.00 0.33 0.43 0.60 0.76 0.93
8………… −93.46 −21.53 −9.30 −4.43 −1.73 −0.57 −0.13 0.02 0.23 0.45 0.67
10………… −117.25 −27.34 −12.05 −5.95 −2.59 −1.13 −0.59 −0.40 −0.13 0.14 0.42
15………… −176.72 −41.85 −18.92 −9.78 −4.73 −2.54 −1.72 −1.45 −1.04 −0.63 −0.22


The velocity of sound in sea water is independent of the wave length except for impulses resulting from the detonation of relatively large amounts of explosives. Initially the impulse from such explosions may travel about 30 per cent faster than normal, apparently because of the tremendous energy involved.

The velocity of sound in a liquid may be computed from the elasticity and density:

but in practice it is more convenient to use the expression
where γ is the ratio of the specific heats, cp/cv (p. 62), ρ is the density, and K is the true compressibility. If ρ and K are in c.g.s. units, the velocity is in centimeters per second. The ratio γ is introduced because the sound impulse is a wave of compression and, hence, heats the water it passes through. The three variables, γ, ρ, and K, all change with temperature, salinity, and pressure, and therefore must be evaluated for any given set of conditions. For example, water of salinity 34.85 ‰ at 30° has a density of 1.021637 at atmospheric pressure (p = 0), and for these conditions γ = 1.0207 and K = 4.196 × 10−11. Hence,

By means of the above formula, suitable tables have been prepared that give the velocity of sound as a function of temperature, salinity, and pressure. The first practical tables were those of Heck and Service (1924) of the U. S. Coast and Geodetic Survey, but these have been superseded by the British Admiralty Tables prepared by Matthews (1927), which, although based on the same original data, are slightly more consistent. The variations in velocity as a function of temperature and salinity are shown in fig. 17. The effect of pressure, and hence depth, is shown in fig. 18. This effect is almost independent of the temperature and salinity, but the curve shown in the figure is actually for salinity 34.85 ‰ and 0°C. This salinity is the standard reference salinity used in these tables, and corrections due to salinity variations are given in the form of anomalies to be added to or subtracted from the standard values.


Velocity of sound in pure water and in sea water at atmospheric pressure as a function of temperature and salinity.


Effect of pressure upon the velocity of sound in sea water of salinity 34.85 ‰ at 0°.

If the velocity of the sound is known, it is possible to determine the wavelength of sound of different frequencies from the equation λ = υ/n, where λ is the wave length, υ is the velocity, and n is the frequency (table 20). Frequencies between about 25 and 10,000 vibrations per second are easily detected by the human ear, and the maximum sensitivity is at about 1000 per second. Ultrasonic waves of frequencies above those audible to the human ear have certain desirable properties that make them valuable in submarine acoustics, but, as will be shown later, their effective range is much less because their absorption is greater. The effective range has been increased by constructing sound emitters which for small wave lengths give a directional beam, somewhat

analogous to that of a searchlight. In practice the spread is about 10° to 15°. A directional beam serves not only to measure the distance to some reflecting body, but also to determine the direction to the body.

Air Sea water Sea water
ϑ = 20°, S = 34.85 ‰ S = 34.85 ‰
Frequency p = 1 atm ϑ = 0°, p = 1 atm ϑ = 20°, p = 1 atm
v = 346 rn/sec v = 1445.2 m/sec v = 1518.5 m/sec
wave length, λ (m) wave length, λ (m) wave length, λ (m)
10 35.6 144.5 151.8
100 3.56 14.45 15.18
1,000 0.36 1.44 1.52
10,000 0.036 0.144 0.152
40,000 0.009 0.036 0.038
100,000 0.0036 0.0144 0.0152

In the absence of appreciable absorption and refraction the intensity of sound varies inversely as the square of the distance from the source. Owing to the viscosity of the water, a certain amount of the kinetic energy of the sound waves is converted into heat, and hence there is an absorption of sound somewhat analogous to that of light. The problem of absorption of sound in sea water has been discussed by Langevin (1924). The sound intensity of a plane sound wave decreases exponentially from I0, to Ix, by passage through a distance x. Therefore

where ν = 8π2μ/3λ2ρυ and corresponds to the absorption coefficient for radiant energy. The distance d over which the intensity is reduced to 1 /e (approximately 1/3) is therefore
where all values are in c.g.s. units. The quantities λ, ρ, and υ have already been defined, and μ is the dynamic viscosity. The ranges of ρ and v are small, and therefore d varies as λ2/μ. Hence the absorption increases rapidly with increasing frequency and somewhat with increasing viscosity, and is significant only for ultrasonic waves. According to Bergmann
(1939) the absorption in water is much greater than is indicated by the equations above. Hartmann and Focke (1940) have obtained experimental data which indicate that the absorption is approximately a thousandfold larger. Whether absorption in the sea is as great or greater than that shown by these laboratory tests must await investigation in the field.

The velocity with which a sound wave travels through the water varies considerably with depth. Hence a “beam” of ultrasonic waves may be refracted when transmitted in a quasi-horizontal direction. Generally the velocity decreases with increasing depth in the upper layers and the beam is bent downward. Studies made by the U. S. Coast and Geodetic Survey (Swainson, 1936) have shown that a sound impulse may travel directly to the hydrophone, or may reach the hydrophone after having been reflected one or more times at the surface and the bottom. In many cases it was possible to distinguish several different “rays” which were received after different intervals. The direct transmission could be obtained only at distances less than 20 km and when the depth to the bottom was about 2000 m. The velocity of the direct wave impulse corresponded to that computed from the temperature and salinity, but those wave impulses which were reflected showed an apparent “horizontal” velocity less than the theoretical because of the longer paths traveled. This empirical horizontal velocity depends upon the distance between the vessel and the hydrophone, the depth to the bottom, the bottom profile, the physical properties of the water, and so forth.


Vertical distributions of temperature and salinity off southern California at 32°57′N and 122°07′W, the computed velocity of sound, and the mean sounding velocity.

As stated previously, the vertical velocity is a function of the depth and the distribution of temperature and salinity. Most sonic depth-finding instruments are adjusted for a constant “sounding velocity,” usually 800 to 820 fathoms per second (1463 to 1500 m/sec). In certain cases it is desirable to correct the readings to true depths. This can be done if the distributions of temperature and salinity are known. For different areas of the oceans and for different depths, the British Admiralty Tables contain “mean sounding velocities”; that is, mean velocities from the surface to the stated depth. In general, these first decrease somewhat with depth, because the effect of decreasing temperature dominates, but at greater depths they increase again as the pressure

effect becomes dominant. In fig. 19 are shown the vertical temperature and salinity distributions at a station off the coast of southern California together with the corresponding velocity of sound at all depths as computed from the British Admiralty Tables. The mean sounding velocity is also shown. The latter decreased from 1503 m/sec at the surface to a minimum of 1484 m/sec for depths between 800 and 1800 m, and then increased again to 1496 m/sec at 4000 m.

Absorption of Radiation

Absorption Coefficients of Distilled Water and of Pure Sea Water. In water the intensity of parallel beams of radiation of wave length λ decreases in the direction of the beams, the decrease in a layer of infinitesimal thickness being proportional to the energy, I, and to the thickness of the layer:

The coefficient of proportionality, x, is called the absorption coefficient. Integrating this equation between the limits x = h and x = h + L, one obtains
where the factor 2.30 enters because base-10 logarithms are used instead of natural logarithms, and where L is the thickness of the layer within which the energy of the radiation is reduced from Iλ,h to Iλ,(h+L) The latter equation also serves as a definition of the absorption coefficient. The numerical value of the absorption coefficient depends upon the unit of length which L is expressed. In physics the unit is 1 cm, but in oceanography it has become common practice to use 1 m as the unit of length. Therefore the numerical values of the coefficients that will be given here are one hundred times larger than the corresponding values given in textbooks of physics.

The decrease of intensity of radiation passing through a layer of water depends not only upon the amount of radiation that is truly absorbed—that is, converted into another form of energy—but also upon the amount that is scattered laterally. In “pure” water the scattering takes place against the water molecules, and the effect of scattering is related to the molecular structure of the water (p. 47). However, when the absorption in pure water is measured, the effect of scattering is not separated but is included in the absorption coefficient, which varies greatly with wave length.

A great number of measurements of the absorption coefficients in pure water have been conducted, but the results by different investigators do not agree (Dorsey, 1940). Thus, at a wave length of .48 μ

(1 μ = 0.0001 cm), at which the absorption is very small, the following values have been reported: Owing to such discrepancies the absorption in pure water is not exactly known, but, as a basis for comparison, table 21 contains values of the absorption coefficients according to determinations by W. R. Sawyer in the range 0.35 μ to 0.65 μ, and by J. R. Collins for wave lengths greater than 0.65 μ (Dietrich, 1939). Sawyer's results have been selected because Clarke and James (1939) have obtained similar values in their examination of filtered sea water.

Hüfner and Albrecht, 1891.................................0.048
Ewan, 1895...............................................0.030
Aschkinass, 1895..........................................0.020
Sawyer, 1931.............................................0.015

From the table it is evident that water is very transparent for radiation of wave lengths between 0.4 μ and 0.6 μ; that is, for visible rays in the violet, blue, green, and yellow parts of the spectrum. It is less transparent for orange and red light, and in the infrared the transparency is practically nil (fig. 21), because, if the absorption coefficient per meter equals 100, 99.5 per cent of the radiation is absorbed in a layer of thickness 5.3 cm.

Wave length in μ Absorption coefficient per meter Wave length in μ Absorption coefficient per meter Wave length in μ Absorption coefficient per meter Wave length in μ Absorption coefficient per meter
.32 0.58 .52 0.019 .85 4.12 1.60 800
.34 0.38 .54 0.024 .90 6.55 1.70 730
.36 0.28 .56 0.030 .95 28.8 1.80 1700
.38 0.148 .58 0.055 1.00 39.7 1.90 7300
.40 0.072 .60 0.125 1.05 17.7 2.00 8500
.42 0.041 .62 0.178 1.10 20.3 2.10 3900
.44 0.023 .65 0.210 1.20 123.2 2.20 2100
.46 0.015 .70 0.84 1.30 150 2.30 2400
.48 0.015 .75 2.72 1.40 1600 2.40 4200
.50 0.016 .80 2.40 1.60 1940 2.50 8500

Collins has compared the absorption in distilled water to that in salt solutions, and from his results it can be concluded that the dissolved salts in concentrations occurring in sea water exert a negligible effect on the absorption coefficient. The maximum effect appears to be about 1.3 per cent, and the uncertainty of the observed values is greater than

this amount. These results have been confirmed by Clarke and James (1939), who found that the absorption of pure sea water as represented by Berkefeld-filtered oceanic sea water was practically identical with that of distilled water. Their observations indicate that Sawyer's values for distilled water may be too high in the ultraviolet; that is, at wave lengths smaller than 0.38 μ.

It has also been concluded that the effect of temperature on absorption, which has been established in the case of distilled water, is applicable to uncontaminated sea water. The effect of temperature changes is to increase the absorption in certain parts of the infrared by about 0.5 per cent for every temperature increase of 1°C, but over a large part of the spectrum the temperature effect is much smaller. When dealing with sea water the effect can be neglected.

Extinction Coefficients in the Sea. In oceanography the greater interest is attached to the rate at which downward-traveling radiation decreases. The rate of decrease can be defined by means of a coefficient similar to the absorption coefficient:

where Iλ,s and Iλ,(z+1) represent the radiation intensities of wave length λ on horizontal surfaces at the depths z and (z + 1) m. Different names have been proposed for this coefficient, such as transmissive exponent (Clarke, 1933) or extinction coefficient (Pettersson, 1936a). The latter name has been widely used and will be employed here, although the process by which the intensity of radiation is reduced will be called absorption. The absorption of radiation in the sea is complicated by the increased scattering due to suspended particles and by the presence of dissolved colored substances. The extinction coefficient of radiation of a given wave length therefore varies within wide limits from one locality to another, and in a given locality it varies with depth and time.

The first crude measurements of absorption in the visible part of the spectrum were made by lowering a white disc of standard size (30 cm), the Secchi disc, and observing the depth at which the disc disappeared from sight. Comparisons with recent exact measurements by other methods have shown that in the English Channel the extinction coefficient of visible rays can roughly be attained from the formula κ = 1.7/D, where D is the maximum depth of visibility in meters, as determined by the Secchi disc (Poole and Atkins, 1929).

The next step in the investigation of the absorption of radiation in sea water was made by subsurface exposure of photographic plates enclosed in watertight containers. Such experiments, which were conducted by Helland-Hansen (1912a) on the Michael Sars Expedition by exposing panchromatic plates at different depths in the vicinity of the Azores, showed that photographic plates were blackened at very great depths.

A plate exposed for 40 minutes at a depth of 500 m showed strong blackening, another exposed for 80 minutes at 1000 m was also blackened, but a third plate which was exposed for 120 minutes at 1700 m showed no effect whatever. These experiments were made at noon on June 6, 1910, with a clear sky. At 500 m it was found that the radiation had a distinct downward direction, because plates exposed at the top of a cube were much more strongly blackened than those exposed on the sides.

In other experiments colored filters were used, which showed that the red portion of the spectrum was rapidly absorbed, whereas the green and blue rays penetrated to much greater depths. Quantitative results as to the absorption at different wave lengths were obtained by using spectro- photometers (Knudsen, 1922), but the methods were laborious and not sensitive enough to be used at great depths.

The introduction in recent years of photoelectric cells has made possible rapid and accurate determinations of extinction coefficients in different parts of the spectrum. A number of different instruments have been and still are in use, but a standardized technique has been proposed by a committee of the International Council for the Exploration of the Sea (Atkins et al, 1938). Because of the wide variation in absorption at different wave lengths, efforts have been directed toward measuring exactly the absorption in narrow spectral bands. The determinations are accomplished by lowering stepwise a photoelectric or photronic cell enclosed in a watertight container and provided with suitably colored filters, and by observing on deck the photoelectric current by means of a sensitive galvanometer or a suitable bridge circuit. The measurements must be made at constant incident light either on clear, sunny days or on days when the sky is uniformly overcast, because the rapid variations in incident light that occur on days with scattered clouds will naturally lead to erroneous results as to the absorption. To determine the percentage amount of radiation that reaches a certain depth, it is necessary to make simultaneous readings of the incident radiation on board ship. For the different precautions that have to be taken, reference is made to papers listed in the bibliography, particularly to Atkins et al (1938).

These methods give information as to the absorption in layers of definite thickness. Instruments for measurements of the transparency of sea water at given depths and of the scattering of light have been designed by H. Pettersson (1936b) and have been used for determining relative values. It has been demonstrated, particularly, that at boundary surfaces sharp variations in transparency and scattering occur. The study of the absorption of radiation in the sea is in rapid progress, and several of the following generalizations are therefore presented with reservations.

The main results as to the character of the extinction coefficient in the sea of radiation of different wave lengths can be well illustrated

by means of data which Utterback (1936) and Jorgensen and Utterback (1939) have published. Utterback attempted to determine the extinction coefficients within spectral bands as narrow as possible, and has assigned the observed coefficients to distinct wave lengths, but it should be understood that the wave length actually stands for a spectral band of definite width. He has made numerous observations in the shallow waters near islands in the inner part of Juan de Fuca Strait and at four stations in the open oceanic waters off the coast of Washington, and these can be considered typical of coastal and oceanic water, respectively. Table 22 contains the absorption coefficients of pure water at the wave lengths used by Utterback, the minimum, average, and maximum extinction coefficients observed in oceanic water, and the minimum, average, and maximum coefficients observed in coastal water. The minimum and maximum coefficients have all been computed from the four lowest and the four highest values in each group. The data in table 22 are represented in fig. 20. In the clearest oceanic water the extinction coefficients were only twice those of pure water and the average values were four to five times the latter, whereas the maximum values were up to ten times as great. In the coastal waters the minimum values were up to sixteen times greater than the absorption coefficients of pure water, the average values were up to twenty-four times as great, and the maximum values were up to thirty-four times as great. The increase of the extinction coefficients, however, varied widely in the different parts of the spectrum and was much greater for shorter wave lengths than for longer.

Type of water Wave length in μ
.46 .48 .515 .53 .565 .60 .66 .80 1.00
Pure water .015 .015 .018 .021 .033 .125 .280 2.40 39.7
Oceanic water lowest .038 .026 .035 .038 .074 .199
Oceanic water average 086 .076 .078 .084 .l08 .272
Oceanic water highest .160 .154 .143 .140 .167 .333
Coastal water lowest .224 .230 .192 .169 .375 .477
Coastal water average .362 .334 .276 .269 .437 .623
Coastal water highest .510 .454 .398 .348 .489 .760


The transparency of the water for radiation of different wave lengths can be expressed by means of the percentage amounts of radiation which penetrate a 1-m layer. These percentage amounts are given in table 23, from which it is seen that the greatest transparency of the clearest oceanic water is at a wave length of 0.48 μ—that is, in the blue part of the spectrum—whereas the greatest transparency of coastal water is at wave lengths 0.53 μ or higher—that is, in the green or green-yellow part of the spectrum. It is also seen that 97.5 per cent of radiation of wave length 0.48 μ passes through 1 m of the clearest oceanic water, but only 63.5 per cent of radiation of the same wave length passes through 1 m of turbid coastal water.


Extinction coefficients of radiation of different wave lengths in pure water and in different types of sea water.

The great difference between the mean and the maximum values of the extinction coefficients shows that the absorption of sea water varies within very wide limits. In the example presented in table 22 the percentage variations are about the same in coastal and oceanic water, and the maximum values in the oceanic water approach the minimum values in the coastal water. In any given locality, great variations also occur in a vertical direction, layers of low absorption alternating with layers of high absorption, and this feature further complicates the actual conditions.

Similar results have been obtained by other investigators from such widely different areas as in the English Channel (Poole and Atkins, 1929;

Poole, 1936), in the waters off the east coast of the United States (Clarke, 1933), and off southern California (Young and Gordon, 1939). In all instances it has been found that the absorption is less in oceanic than in coastal water, but it varies within wide limits both locally and with depth. Where examination of absorption in different parts of the spectrum has been conducted, it has been found that the absorption is much less in the blue than in the red end of the spectrum and that the blue light penetrates to the greatest depths in clear water, whereas the green or yellow light reaches further down in turbid water.

Type of water Wave length (μ)
.46 .48 .515 .53 .565 .60 .66
Pure water 98.5 98.5 98.2 97.9 96.8 88.3 75.9
Oceanic water lowest 96.4 97.5 96.6 96.3 92.9 81.8
Oceanic water average 91.8 92.7 92.5 91.8 89.8 75.9
Oceanic water highest 85.1 85.7 86.7 86.9 84.5 71.6
Coastal water lowest 80.0 79.4 82.6 84.5 68.7 62.0
Coastal water average 69.7 71.6 75.9 76.4 64.6 53.6
Coastal water highest 60.0 63.5 67.1 70.6 61.4 46.7

Influence of the Altitude of the Sun Upon the Extinction Coefficient. The extinction coefficient is a measure of the reduction of intensity on a, vertical distance and depends, therefore, upon the obliquity of the rays. The obliquity of the incident rays is reduced, however, by refraction when entering the water from the air and by the effect of scattering. When the sun's rays pass the water surface, the angle of refraction increases from zero with the sun in zenith to 48.5 degrees with the sun at the horizon, and therefore the most oblique rays penetrating into the water form an angle of less than 48 degrees with the vertical. Owing to the scattering and the sifting out by absorption of the most oblique rays, the radiation that penetrates to moderate depths will become nearly vertical, and the measured extinction coefficients will be independent, within wide limits, of the altitude of the sun. The reduction of the obliquity of the incident radiation has been directly demonstrated by Johnson and Liljequist (1938). Conditions at very. low sun have not been examined, but it is probable that at low sun the extinction coefficients are increased, and this increase may have bearing

upon the diurnal variation of the incoming energy at greater depths (p. 779).

The Scattering of Radiation in the Sea The Scattering of Radiation has been examined both directly by means of Pettersson's scattering meter (p. 83) and by measuring the relative intensities of downward- and upward-traveling radiation or vertical and horizontal radiation. Jorgensen and Utterback (1939) found that in coastal waters the intensity of the upward-traveling radiation ranged for the short wave lengths from 1 to 3 per cent of that of the downward-traveling radiation, and for the long wave lengths from 0.5 to 2 per cent. In oceanic water Utterback (1936) found ratios between 1 and 2 per cent at the shorter wave lengths. Clarke (1936) found considerably higher values in shallow coastal waters, but similar values in the deep basin of the Gulf of Maine.

The relative intensities of horizontal and vertical radiation have been measured by Clarke off the east coast of the United States and by Poole and Atkins in the English Channel. The greatest value found by Clarke was 17 per cent, but Poole and Atkins (1929) have reported an average value of 50 per cent for the horizontal radiation down to a depth of 25 m in the English Channel. The conclusion that can be drawn from these experiments is that the subsurface illumination becomes more and more diffuse with increasing depth, particularly in coastal waters, but that the directional character of the radiation is lost only slowly. This conclusion is particularly true in clear oceanic water, where Helland-Hansen (1912a) found that the vertical radiation was distinctly more intense than the horizontal at a depth of 500 m. (p. 82).

Cause of the Great Extinction Coefficients in the Sea. The great extinction coefficients in the sea as compared to those of absolutely pure water are as a rule ascribed to the presence of minute particles which cause scattering and reflection of the radiation and which themselves absorb radiation. If such particles are small compared to the wave length, λ, of the radiation, the scattering will be proportional, according to Lord Rayleigh, to λ−4, and the effect therefore at wave length, say, .46 μ will be 2.86 times greater than at wave length .60 μ. This selective effect leads to a shift toward longer wave lengths in the region of minimum absorption.

Clarke and James (1939) found that the increased absorption in oceanic water was chiefly caused by suspensoids that could be removed by means of a “fine” Berkefeld filter, and that these suspensoids were largely nonselective in their effect. Utterback's data (1936) indicated, on the other hand, that the increased absorption in oceanic water is at least in part due to selective scattering, because at short wave lengths the extinction coefficients were increased more, above those of pure water, than at longer wave lengths (table 22). Kalle (1938) is of the

opinion that selective scattering is of dominant importance (p. 89), but the question is not yet settled as to the mechanism which leads to increased absorption in oceanic water as compared to pure water. The fact that even in the clearest oceanic water the absorption is greater than in pure water indicates, however, that finely suspended matter is always present. One could state that the ocean waters always contain dust.

The increase of the absorption coefficients in coastal waters appears to be due in part to another process. Clarke and James (1939) conclude from their examination that in coastal water both suspensoid and “filter-passing” materials are effective in increasing the absorption, and that each exerts a highly selective action, with greatest absorption at the shorter wave lengths. These great absorptions at the shorter wave lengths are demonstrated by Utterback's measurements (table 22). Clarke does not discuss the nature of the “ filter-passing” material, but Kalle (1938) has shown that in sea water water-soluble pigments of yellow color are present. These pigments appear to be related to the humic acids, but their chemical composition has not been thoroughly examined, for which reason Kalle calls them “yellow substance.” This yellow substance seems to occur in greatest abundance in coastal areas, but Kalle has demonstrated its presence in the open ocean as well and believes that it represents a fairly stable metabolic product related to the phytoplankton of the sea. The selective absorption 'of this yellow substance may then be responsible, in part, for the character of the absorption in coastal water and for shift of the band of minimum absorption toward longer wave lengths.

It has not been possible anywhere to demonstrate any direct influence of phytoplankton populations on the absorption, but very dense populations may cut down the transparency. At present it appears that the major increase of absorption of sea water over that of pure water is due to two factors: the presence of minute suspended particles, and the presence of dissolved “yellow substance.” The former factor dominates in oceanic waters, and the latter is particularly important in coastal waters.

The Color of the Sea. The color of the sea, as it appears to an observer ashore or on board a vessel, varies from a deep blue to an intense green, and is in certain circumstances brown or brown-red. The blue waters are typical of the open oceans, particularly in middle and lower latitudes, whereas the green water is more common in coastal areas, and the brown or “red” water is observed in coastal regions only.

The color of the sea has been examined by means of a Secchi disk (p. 82) by observing the color that the water appears to have when seen against the white submerged surface of the disc. This color is recorded according to a specially prepared color scale, the “ Forel scale” (Krümmel, 1907). The method is a rough one and the scale is not adapted for

recording the extreme colors in coastal waters. In order to obtain more exact results, Granquist, working in the Finnish waters, used a long tube that was blackened on the inside, and this type of instrument has been greatly improved by Kalle (1938).

Kalle has critically reviewed earlier theories as to the causes of the color of the sea and arrives at conclusions that appear to be consistent with all available observations. The blue color is explained, in agreement with earlier theories, as a result of scattering against the water molecules themselves, or against suspended minute particles smaller than the shortest visible wave lengths. The blue color of the water is therefore comparable to the blue color of the sky. The transition from blue to green cannot be explained, however, as a result of scattering, and Kalle concludes that this transition is due to the above-mentioned “ yellow substance,” pointing out that the combination of the yellow color and the “natural” blue of the water leads to a scale of green colors as observed at sea. Fluorescence may contribute to the coloring but appears to be of minor importance.

The color of suspended larger particles, if present in great abundance, can give color to the sea. In this case the color is not determined by the optical properties of the water or by dissolved matter, but by the colors of the suspended inorganic or organic particles, and the water is appropriately called “ discolored.” Discoloration can be observed when large quantities of finely suspended mineral particles are carried into the sea after heavy rainfall, or when very large populations, several million cells per liter, of certain species of algae or dinoflagellates are present very near the surface. Thus, the “red water” (often more brown than red) which is quite frequently observed in many areas and after which the Red Sea and the Vermilion Sea (Gulf of California) have been named is due to abundance of certain algae (in the Red Sea, Trichodesmium erythraeum) or dinoflagellates. The discoloration, beautiful examples of which have been given by Gunther (1936), is, however, a phenomenon of the typical coastal waters, the green colors being frequent in waters near the coast or at sea in high latitudes, and the blues charac teristic of the open ocean in middle and lower latitudes (fig. 214, p, 784).

Eddy Conductivity, Diffusivity, and Viscosity

In the preceding discussion it has been repeatedly stated that the coefficients of heat conduction, diffusion, and viscosity that have been dealt with so far are applicable only if the water is at rest or in laminar flow. By laminar flow is understood a state in which sheets (laminae) of liquids move in an orderly manner such that random local fluctuations of velocity do not occur. However, the molecules of the liquid, including those of dissolved substances, move at random, and, owing to this

random motion, an exchange of molecules takes place between adjacent layers. Consequently, there is a transfer of heat if adjacent layers are at different temperatures, a diffusion of dissolved substances if their concentrations are variable in space, or a transfer of momentum if their velocities differ. The amounts that are transferred are proportional to the gradients of temperature, concentration, or velocity, and the factors of proportionality—that is, the coefficients of thermal conductivity, diffusion, and viscosity—are among the characteristic properties of the liquid. For a given liquid they are functions of temperature, pressure, and concentration, and can be exactly determined in the laboratory.

In nature, laminar flow is rarely or never encountered, but, instead, turbulent flow, or turbulence, prevails. By turbulent flow is understood a state in which random motion of smaller or larger masses of the fluid is superimposed upon some simple pattern of flow. The character of the turbulence depends upon a number of factors, such as the average velocity of the flow, the average velocity gradients, and the boundaries of the system. Under these conditions the exchange between adjacent moving layers is not limited to the interchange of molecules, but masses of different dimensions also pass from one layer to another, carrying with them their characteristic properties. As a consequence, a snapshot of the instantaneous distribution of velocity, temperature, salinity, and other variables in the sea would show a most complicated pattern, but so far no means have been developed for establishing this picture. Measurements by sensitive current meters have demonstrated that in a given locality the velocity fluctuates from second to second, but in most cases observations of ocean currents give information as to mean velocities for time intervals that may vary from a few minutes to twenty-four hours or more. Similarly, special measurements have demonstrated that the details of the temperature distribution are very complicated, but in general observations are made at such great distances apart that only the major features of the temperature distribution are obtained. Inasmuch as it is impossible to observe the instantaneous distribution in space of temperature, salinity, and velocity, it follows that the corresponding gradients cannot be determined and that no basis exists for application to the processes in the sea of the coefficients of thermal conductivity, diffusion, and viscosity that have been determined in the laboratory. Since only certain average gradients can be determined, another approach has to be made when dealing with the processes in the sea. In order to illustrate this approach, let us first consider the viscosity.

In the case of laminar flow the coefficient of viscosity, μ, is defined by the equation τs = μdv/dn where τs is the shearing stress exerted on a surface of unit area, and dv/dn is the shear normal to that surface.

In the case of turbulent flow a coefficient of eddy viscosity, A, can be defined in a similar manner:
where dū/dn now represents the shear of the observed velocities. The numerical value of the eddy viscosity depends upon the size and intensity of the eddies—that is, on the magnitude of the exchange of masses between adjacent layers; and the symbol A that is commonly used is an abbreviation for the term “ Austausch” introduced by Schmidt (1917). The numerical value of A also depends upon how the “average” velocities have been determined; that is, upon the distribution in space of the observations and upon the length of the time intervals to which the averages refer.

The definition of the eddy viscosity in the above manner appears purely formalistic, but it is based on the concept that masses which leave one layer carry with them the momentum corresponding to the average velocity in that layer, and that by impact they attain the momentum corresponding to the average velocity of their new surroundings before again leaving them (p. 472). Thus, A is an expression for the transfer of momentum of mean motion. This transfer is much increased by the turbulence, as is evident from the fact that the eddy viscosity is many times greater than the molecular viscosity.

The eddy viscosity can be determined only by examination of the effect on the mean motion. This effect is discussed on pp. 492 and 577, but a few points will be mentioned here. It has been found practical to distinguish between two types of turbulence in the sea—vertical and lateral, In the case of vertical turbulence the effective exchange of masses is related to comparatively slight random motion in a vertical direction or, if the term “eddy motion” is used, to small eddies in a vertical plane. Actually, the eddies are oriented at random, but only their vertical components produce any effect on the mean motion. The corresponding eddy viscosity has been found to vary between 1 and 1000 c.g.s. units, thus being one thousand to one million times greater than the molecular viscosity of water. In the case of lateral turbulence the effective exchange of masses is due to the existence of large quasi-horizontal eddies, The corresponding eddy viscosity depends upon dimensions of the system under consideration and has been found to vary between 106 and 108 c.g.s. units.

The distinction between vertical and lateral turbulence is particularly significant where the density of the water increases with depth, because such an increase influences the two types of turbulence in a different manner. Where the density of the sea water increases with depth (disregarding the effect of pressure), vertical random motion is impeded by

Archimedean forces, because a mass which is brought to a higher level will be surrounded by water of less density and will tend to sink back to the level from which it came, and, similarly, a water mass moving downward will be surrounded by denser water and will tend to rise. In this case the stratification of the water is called stable, because it cannot be altered unless work against gravity is performed. Stable stratification reduces the vertical turbulence; where the stability is very great the vertical turbulence may become nearly suppressed and the eddy viscosity small. The effect of stability on the lateral turbulence, on the other hand, is negligible, because the lateral random motion takes place mainly along surfaces of equal density.

With regard to the eddy conductivity, similar reasoning is applicable. When dealing with eddy viscosity it was assumed that the exchange of mass leads to a transfer of momentum from one layer to another, which is expressed by means of A. Correspondingly, when dealing with eddy conductivity, one can assume that the transfer of heat through any surface is proportional to the exchange of mass through the surface, as expressed by A, and to the gradient of the observed temperatures, −dformula/dn; that is, dQ/dt = −rA dformula/dn, where r is a factor that depends upon the specific heat of the fluid and upon the manner in which the heat contents of the moving masses are given off to the surroundings. When dealing with homogeneous water, it is assumed that a large mass which is transferred to a new level breaks down at that level into smaller and smaller elements, and that equalization of temperature ultimately takes place by molecular heat conduction between the small elements and the surroundings. If such is the case, both the difference in momentum and the difference in heat content are leveled off, and the proportionality factor, r, is equal to the specific heat of the liquid. Since the specific heat of water is nearly unity, the numerical values of eddy conductivity and eddy viscosity are practically equal. However, where stable stratification prevails, the elements, being lighter or heavier than their surroundings, may return to their original level before completion of temperature equalization, but equalization of momentum may have been accomplished by collision. In this case, the factor of proportionality, r, will be smaller than the specific heat of the liquid; that is, in the sea, r is smaller than unity, and the eddy conductivity is smaller than the eddy viscosity. Thus, stable stratification reduces the vertical eddy conductivity even more than it reduces the vertical eddy viscosity. Taylor (1931) has presented the above reasoning in mathematical language (p. 476).

The discussion has so far been limited to a consideration of the vertical eddy conductivity, but lateral eddy conductivity due to lateral turbulence has also to be introduced. The numerical value of lateral

eddy conductivity must be nearly equal to that of the lateral eddy viscosity, because the lateral turbulence is not affected by stable stratification.

Numerical values of the coefficients of eddy conductivity can be derived only from a study of the effect of mixing processes on the observed distribution of temperature. Methods of such determinations and numerical values are presented on pp. 483 and 484. The results have confirmed the above conclusions and have also demonstrated that the eddy conductivity varies within wide limits. In the upper layers of the sea, where stable stratification prevails, the vertical coefficient of eddy viscosity varies between 1 and 1000, whereas the corresponding eddy conductivity is smaller and varies between 0.01 and 100; in homogeneous water, however, no difference has been established (p. 485). In the cases in which the lateral coefficients of viscosity and conductivity have been examined, nearly equal numerical values have been found in agreement with the conclusion that the stability of the stratification does not influence the lateral turbulence.

The transfer of salinity or other concentration is similar to the heat transfer. The eddy diffusivity is also proportional to the exchange of mass as expressed by A, the factor of proportionality being a pure number. In sea water of uniform density r = 1, but in the case of stable stratification, when complete equalization of concentration does not take place, r < 1; that is, the vertical eddy diffusivity is smaller than A and equals the eddy conductivity. This conclusion has also been confirmed by observations (p. 484).


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Thomas, B. D., T. G. Thompson, and C, L. Utterback. 1934. The electrical conductivity of sea water. Conseil Perm. Intern, p. l'Explor. de la Mer, Jour, du Conseil, v. 9, p. 28–35, 1934.

Thompson, T. G.1932. “The physical properties of sea water” . Physics of the earth, v. 5, Oceanography, p. 63–94. Nat. Research Council, Bull. no. 85, 1932.


Utterback, C. L.1936, Spectral bands of submarine solar radiation in the North Pacific and adjacent inshore waters. Conseil Perm. Intern, p. l'Explor. de la Mer, Rapp. et Proc.-Verb., v. 101, pt. 2, no. 4, 15 pp., 1936.

Utterback, C. L., T. G. Thompson, and B. D. Thomas. 1934. Refractivity-chlorinity-temperature relationships of ocean waters. Conseil Perm. Intern, p. l'Explor. de la Mer, Jour, du Conseil, v. 9, p. 35–38, 1934.

Wenner, P., Edward H. Smith, and Floyd M. Soule. 1930. “Apparatus for the determination aboard ship of the salinity of sea water by the electrical conductivity method” . Jour. Research, U. S. Bur. Standards, v. 5, p. 711–32, 1930.

Wirth, H. E., T. G. Thompson, and C. L. Utterback. 1935. “Distribution of isotopic water in the sea” . Amer. Chem. Soc., Jour,, v. 57, p. 400–04, 1935.

Witting, R.1908. “Untersuchungen zur Kenntnis der Wasserbewegungen und der Wasserumsetzung in den Finnland umgebenden Meeren” . Finländische Hydrogr.-Biologische Untersuchungen, no. 2, 246 pp., 1908. Cf. p. 173.

Young, R. T., Jr., and R. D. Gordon. 1939. “Report on the penetration of light in the Pacific Ocean off the coast of southern California” . Scripps Inst. Oceanogr., Univ. Calif., Bull., tech. ser., v. 4, p. 197–218, 1939.


IV. General Distribution of Temperature, Salinity, and Density

The Heat Budget of the Earth as a Whole

For the earth as a whole, the total amount of heat that is received during one year from the sun at the limit of the atmosphere must exactly balance the total amount that in the same period is lost by reflection and by radiation into space. Otherwise, the temperature of the atmosphere and the oceans would change. The radiation from the hot sun is called short-wave radiation, because the wave lengths which reach the limit of the earth's atmosphere lie roughly between 0.38 μ and 2.5 μ, whereas the dark-heat radiation which is emitted by all objects at ordinary temperatures is called long-wave radiation, being of wave lengths between 5 μ and 20 μ. The part of the short-wave radiation that is reflected is of no importance to the heat budget of the earth, and therefore the amount of short-wave radiation that is absorbed by the atmosphere, the oceans, and the land must exactly balance the long-wave radiation into space from the entire system. A small part of the heat that the atmosphere receives is transformed into kinetic energy which by friction is transformed back again to heat and ultimately lost into space by radiation. Thus, the transformation of heat to kinetic energy does not lead to any net gain of heat but serves to maintain the circulations of the atmosphere and the oceans.

As is customary procedure, the amounts of heat will be given in gram calories and not in units of work such as ergs or joules. The conversion factors are: 1 gram calorie = 4.183 × 107 ergs = 4.183 joules.

In lower latitudes, heat received by radiation is greater than heat lost by back radiation and reflection, whereas in higher latitudes the gain is less than the loss. Table 24 contains values of heat received and lost by processes of radiation and reflection in different latitudes. The third column, containing the differences between the two quantities, shows that there is an annual net gain of heat in the equatorial regions and a net loss in the polar regions. The mean annual temperatures in different latitudes on the earth remain unchanged from one year to another, showing that within the atmosphere and the oceans there must be a transport of heat from lower to higher latitudes which exactly equals the difference

between heat received and heat lost by radiation. Multiplying the average difference between any two parallels of latitude by the area of the earth's surface between these parallels and summing up, starting at the Equator, gives the total amounts of heat that flow from the Equator toward the poles in every latitude. Some of these values are given in the fourth column in the table, from which it is seen that they are of the order of 1016 g cal/min. Dividing the numbers by the length of the parallels gives the amounts shown in the fifth column of the table, which represent the average flow of heat across each centimeter of the parallels of latitude. These numbers are of the order of 107 g cal/cm/min.

The transport of heat from lower to higher latitudes takes place partly by air currents (winds) and partly by ocean currents. In meteorological literature it is generally assumed that the transport by ocean currents is negligible (Bjerknes et al, 1932), although the question has not been thoroughly examined. It can be shown that the assumption is correct when dealing with averages for the whole earth, but in some regions the transport by ocean currents is of considerable importance.

Latitude (°) Heat received (g cal/cm2/min) Heat lost (g cal/cm2/min) Surplus or deficit (g cal/cm2/min) Heat transport across parallels of latitude (1016 g cal/min) Heat transport across every centimeter of parallels of latitude (107 g cal/cm/min)
0 0.339 0.300 0.039 0.00 0.00
10 0.334 0.299 0.035 1.59 0.40
20 0.320 0.294 0.026 2.94 0.78
30 0.297 0.283 0.014 3.58 1.07
40 0.267 0.272 −0.005 3.96 1.30
50 0.232 0.258 −0.026 3.34 1.32
60 0.193 0.245 −0.052 2.40 1.20
70 0.160 0.231 −0.071 1.20 0.88
80 0.144 0.220 −0.076 0.32 0.46
90 0.140 0.220 −0.080 0.00 0.00

The amount of heat transported in a north-south direction by a unit volume of ocean water is equal to cρϑvN, where c represents the specific heat, ρ the density, ϑ the temperature of the water, and vN the north-south component of velocity. The total transport through a certain section of the sea can be found by integration, but for the sake of simplicity

we shall assume that this transport can be written cformulaρTN, where formula represents the average temperature of the water and ρTN represents the mass of water passing north through the section in unit time. If the section is taken across an ocean, the mass transport to the north, ρTN, must equal the mass transport to the south, ρTS, but the heat transport may differ because the temperature of the water transported in one direction may be higher or lower than that of the water which is transported in the opposite direction. If these temperatures are designated by ϑN and ϑS, the net transport of heat will be cN − ϑST, where ρT now means the transport to the north and the south. As an example, we can apply these considerations to the North Atlantic Ocean along the parallel of 55°N. In the eastern Atlantic about 10 million m3/sec of warm water flow toward the north, and in the western Atlantic an equal volume of cold water is carried south by the Labrador Current and by the flow of the deep water (p. 684). With ϑN − ϑS = 5°, c = 1, ρ = 1, and T = 10 × 106 m3/sec, we find that the heat transport toward the north through latitude 55° in the Atlantic Ocean is about 0.3 × 1016 g cal/min. The total heat transport across the parallel of 55°N is about 3 × 1016 g cal/min; hence in the North Atlantic the fraction carried by the ocean currents is appreciable.

This example represents an outstanding case of poleward transport of heat by ocean currents. In the Pacific Ocean a transport of comparable magnitude probably takes place in latitudes 30°N to 40°N, but in the southern oceans the north-south circulation and the corresponding temperature contrasts between currents flowing toward or away from the higher latitudes are smaller than those in the northern oceans. A detailed study of the transport of heat by ocean currents has not been made, but it seems certain that by far the major transport of heat is taken care of by the atmosphere.

The Heat Budget of the Oceans

The above consideration applies to the entire system formed by the atmosphere and the oceans, but for the oceans alone we encounter an entirely different picture. On an average, the gain of heat must exactly balance the loss, but the processes involved are not limited to those of radiation, as is evident from the list at the bottom of page 101.

These processes will be discussed in detail, but it can already be stated that of the processes of heating only the first one is important, and the heat budget of the oceans as a whole can therefore be written

where Qs is the heat received. Over all ocean surfaces between 70°N and 70°S the average amounts (in g cal/cm2/min) are, according to Mosby (1936),

If a specific region is considered, it must be taken into account that heat may be brought into or out of that region by ocean currents or by processes of mixing, and that during short time intervals a certain amount of heat may be used for changing the temperature of the water. The complete equation for the heat balance of any part of the ocean in a given time interval is, therefore,

where Qv represents the net amount which by currents or processes of mixing is brought into or out of the region, and where Qϑ represents the amount of heat used locally for changing the temperature of the sea water.

Radiation from the Sun and the Sky. The short-wave radiation that reaches the sea surface comes partly directly from the sun and partly from the sky as reflected or scattered radiation. The amount of radiation energy which is absorbed per unit volume in the sea depends upon the amount of energy that reaches the sea surface, the reflection from the sea surface, and the absorption coefficients for total energy. The incoming radiation depends mainly upon the altitude of the sun, the absorption in the atmosphere, and the cloudiness. With a clear sky and a high sun, about 85 per cent of the radiation comes directly from the sun and about 15 per cent from the sky, but with a low sun the proportion from the sky is greater, reaching about 40 per cent of the total with the sun 10 degrees above the horizon.

Processes of Heating of the Ocean Water Processes of Cooling of the Ocean Water
  1. Absorption of radiation from the sun and the sky, Qs.

  2. Convection of heat through the ocean bottom from the interior of the earth.

  3. Transformation of kinetic energy to heat.

  4. Heating due to chemical processes.

  5. Convection of sensible heat from the atmosphere.

  6. Condensation of water vapor.

  1. Back radiation from the sea surface, Qb.

  2. Convection of sensible heat to the atmosphere, Qh.

  3. Evaporation, Qe.


The incoming energy from the sun is cut down when passing through the atmosphere, partly through absorption by water vapor and carbon dioxide in the air and partly through scattering against the air molecules or very fine dust. The total effect of absorption and scattering in the atmosphere depends upon the thickness of the air mass through which the sun's rays pass, as expressed by the equation

Here I represents the energy in g cal/cm2/min reaching a surface which is normal to the sun's rays; m represents the relative thickness of the air mass and is equal to 1 at a pressure of 760 mm when the sun stands in zenith, equal to 2 when the sun is 30° above the horizon (sin 30° = ½), and so on; S is the solar constant (1.932 g cal/cm2/min); T is the “turbidity factor” of the air; and am is 0.128−0.054 log m.

The sun's radiation on a horizontal surface is obtained by multiplication with sin h, where h is the sun's altitude. To this amount must be added the diffuse sky radiation in order to obtain the total radiation on a horizontal surface. Instruments are in use for recording the total radiation and for recording separately the radiation from the sun and from the sky.

When the sun is obscured by clouds, the radiation comes from the sky and the clouds and, on an average, can be represented by the formula Q = Q0(1 − 0.071 C), where the cloudiness C is given on the scale 0 to 10, and where Q0 represents the total incoming radiation with a clear sky. This formula is applicable, however, only to average conditions. If the sun shines through scattered clouds, the radiation may be greater than with a clear sky, owing to the reflection from the clouds, and on a completely overcast dark and rainy day the incoming radiation may be cut down to less than 10 per cent of that on a clear day. Table 25 contains the average monthly amounts of incoming radiation, expressed in g cal/cm2/min, which reach a horizontal surface in the indicated localities (computed from Kimball, 1928). The differences between the parts of the oceans in the same latitudes are mainly due to differences in cloudiness.

Few direct measurements of radiation are available from the oceans, and when dealing with the incoming radiation it is necessary to consider average values which can be computed from empirical formulae. Mosby (1936) has established such a formula by means of which monthly or annual mean values of the incoming radiation on a horizontal surface can be computed if the corresponding average altitude of the sun and the average cloudiness are known:

Here formula is the average altitude of the sun. The factor k depends upon the transparency of the atmosphere and appears to vary somewhat with latitude, being 0.023 at the Equator, 0.024 in lat. 40°, and 0.027 in lat. 70°. Mosby's formula is not valid at h > 60°, but gives correct results if, at high altitudes of the sun, the true altitude is replaced by a reduced altitude as follows: The values computed by means of this formula agree within a few per cent with those derived by Kimball in an entirely different manner (table 25).

Locality Month
Latitude Longitude Jan. Feb. Mar. Apr. May June July Aug. Sept. Oct. Nov. Dec.
60°N 7°E- 56°W .002 .053 .125 .207 .272 .292 .267 .212 .147 .074 .006 0
60 N 135–170 W .005 .078 .155 .208 .269 .260 .242 .185 .127 .077 .015 0
52 N 10 W .048 .089 .148 .219 .258 .267 .251 .211 .160 .104 .062 .041
52 N 129 W .053 .091 .135 .185 .246 .250 .230 .214 .158 .097 .058 .039
42 N 66 -70W .094 .138 .212 .272 .306 .329 .302 .267 .230 .174 .115 .086
42 N 124 W .100 .151 .210 .286 .331 .360 .320 .274 .231 .174 .113 .092
30 N 65–77 W .146 .165 .238 .285 .317 .310 .301 .282 .239 .188 .169 .142
30 N 128 -130 E .141 .153 .199 .241 .258 .238 .256 .260 .219 .178 .153 .135
10 N 61–69 W .254 .276 .299 .305 .272 .276 .285 .292 .287 .269 .248 .239
10 N 116 E- 80 W .226 .257 .292 .278 .255 .239 .240 .242 .247 .237 .224 .219
0 7–12 E .239 .248 .244 .230 .210 .196 .188 .194 .220 .240 .239 .235
0 48 W & 170 E .261 .265 .282 .297 .309 .300 .300 .340 .366 .362 .339 .278
10 S 14 E; 36- 38 W .329 .328 .301 .254 .219 .206 .232 .278 .312 .324 .317 .320
10 S 72 -171 E .290 .308 .315 .289 .266 .253 .269 .306 .332 .313 .301 .303
30 S 17 and 116 E .452 .406 .340 .254 .186 .148 .166 .214 .274 .362 .401 .430
30 S 110 W .380 .330 .260 .209 .162 .130 .145 .176 .237 .321 .340 .390
42 S 73 W; 147 E .343 .297 .223 .154 .104 .085 .092 .135 .187 .264 .310 .348
52 S 58 W .289 .237 .167 .112 .062 .039 .049 .097 .150 .222 .273 .302
60 S 45 W .213 .171 .105 .056 .011 0 .003 .054 .111 .156 .204 .221

True altitude (°) 60 65 70 75 80 85 90
Reduced altitude (°) 60 62 64 66 68 69 70

Part of the incoming radiation is lost by reflection from the sea surface, the loss depending upon the altitude of the sun. When computing the loss, the direct radiation from the sun and the scattered radiation from the sky must be considered separately. With the sun 90°, 60°, 30° and 10° above the horizon, the reflected amounts of the direct solar radiation are, according to Schmidt (1915), 2.0 per cent, 2.1 per cent, 6.0 per cent, and 34.8 per cent respectively. For diffuse radiation from the sky and from clouds Schmidt computes a reflection of 17 per cent. Measurements by Powell and Clarke (1936) gave values on clear days in agreement with the above, but on overcast days when all radiation reaching the sea surface was diffuse, the observed reflection was about 8 per cent. If the fractions of the total radiation from the sun and the sky on a clear day are designated p and q, respectively, and if the corresponding percentages reflected are designated m and n, the percentage of the total incoming radiation that is reflected on a clear day is r = mp + nq. Thus, on an overcast day, when all incoming radiation is diffuse, r = 8 per cent. Table 26 contains approximate values of r at different altitudes of the sun on a clear day.

Altitude of the sun (°) 5 10 20 30 40 50 60 70 80 90
Percentage reflected 40 25 12 6 4 3 3 3 3 3

The values in the table are applicable only if the sea surface is smooth. In the presence of waves the reflection loss at a low sun is somewhat increased and will be of particular importance in high latitudes. The amount of radiation which under stated conditions penetrates the sea surface is obtained by subtracting the reflection loss from the total incoming radiation.

Absorption of Radiation Energy in the Sea. The radiation that penetrates the surface is absorbed in the sea water. The amounts absorbed within given layers of water can be derived by measuring with

a thermopile the energies which reach different depths or by computing these energies by means of known extinction coefficients. Direct measurements of energy have been made in Mediterranean waters only (Vercelli, 1937), but extinction coefficients of radiation of different wave lengths have been determined in many areas (p. 85). For computation of the energy that reaches a given depth, it is necessary to know the intensity of the radiation at different wave lengths; that is, the energy spectrum. The reduction in intensity has to be calculated for each wave length, and the total energy reaching a given depth has to be determined from the energy spectrum by means of integration. The definition of the extinction coefficient for total energy corresponds to the definition of extinction coefficients at given wave lengths (p. 82).


Schematic representation of the energy spectrum of the radiation from the sun and the sky which penetrates the sea surface, and of the energy spectra in pure water at depths of 0.1, 1, 10, and 100 m. Inset: Percentages of total energy and of energy in the visible part of the spectrum reaching different depths.

The spectrum of the energy that penetrates the sea surface is represented approximately by the upper curve in fig. 21, which also shows the energy spectra at different depths in pure water. The total energy at any given depth is proportional to the area enclosed between the base line and the curves showing the energy spectrum. In the inserted diagram the total energy, expressed as percentage of the energy penetrating the surface, as well as the corresponding percentages of the energy in the visible part of the spectrum, is plotted against depth. The figure shows that pure water is transparent for visible radiation only.

For sea water the percentage of the total energy reaching various depths has been computed for the clearest oceanic water, for average

oceanic water, for average coastal water, and for turbid coastal water, using the extinction coefficients shown in fig. 20. The results are presented in table 27. In the clearest offshore water, 62.3 per cent of the incoming energy is absorbed in the first meter. The absorption is often increased in the upper one meter owing to the presence of foam and air bubbles. This increased absorption, when dealing with the penetration of light, is referred to as “surface loss.” If this process is disregarded, the values clearly demonstrate that the greater amount of energy is absorbed very near the sea surface and that the amount which penetrates to any appreciable depth is considerable only when the water is exceptionally clear. At 10 m, 83.9 per cent has been absorbed in the clearest water and 99.55 per cent in the turbid coastal water.


Energy spectra at a depth of 10 m in different types of water. Curves marked 0, 1, 2, 3, and 4 represent energy spectra in pure water, clear oceanic, average oceanic, average coastal, and turbid coastal sea water, respectively. Inset: Energy spectra at a depth of 100 m in clear oceanic water and at 10 m in turbid coastal water.

The absorption of energy is illustrated in fig. 22, which shows the energy spectra in different types of water at a depth of 10 m. At this depth the maximum energy in the clearest water is found in the blue-green portion of the spectrum, whereas in the turbid coastal water the maximum has been displaced toward the greenish-yellow part. This displacement is further illustrated by the inserted curve in the upper right-hand corner of the figure, which shows the energy spectra at 100 m in the clearest water and at 10 m in the most turbid water.

Extinction coefficients of total energy have been computed and are entered in table 27. These extinction coefficients are very high in the upper 1 m but decrease rapidly, at greater depth approaching the minimum extinction coefficients characteristic of the types of water dealt with. The smallest values given in the table can be considered valid at greater depths as well.

Percentage amounts of incident energy Extinction coefficients per meter
Depth Pure sea water Oceanic water Coastal water Interval of depth (m) Pure sea water Oceanic water Coastal water
Clearest Average Average Turbid Clearest Average Average Turbid
0 100 100 100 100 100 0–1 .944 .975 1.080 1.318 1.385
1 38.9 37.7 35.2 26.7 22.8 1–2 .143 .176 .230 .450 .547
2 33.7 31.6 28.0 17.0 13.2 2–5 .062 .095 .159 .351 .452
5 28.0 23.7 17.3 5.95 3.41 5–10 .048 .076 .120 .318 .405
10 22.0 16.1 9.50 1.21 0.449 10–20 .033 .054 .094 .292 .368
20 15.8 9.35 3.72 0.064 0.012 20–50 .024 .042 .083
50 7.64 2.69 0.311 50–100 .018 .036
100 3.04 0.452 0.0057


In fig. 23 the curves marked 0, 1, 2, 3, and 4 represent the percentage amounts of energy which reach different levels between the surface and 10 m, according to the data in table 27. The three curves marked Capri, Trieste, and Venice represent results of measurements in the Mediterranean according to Vercelli (1937), and four other curves represent observed values in lakes according to Birge and Juday (1929). The agreement of the character of the curves indicates that reliable values as to the absorption of energy in the sea can be obtained by means of computations based on observed extinction coefficients.


Percentages of total energy reaching different depths in pure water, clear oceanic, average oceanic, average coastal, and turbid coastal sea water (curves 0, 1, 2, 3, and 4), computed from extinction coefficients, and corresponding directly to observed values in four lakes and at three localities in the Mediterranean.

An idea of the heating due to absorption of radiation can be secured by computing the increase of temperature at different depths which results from a penetration of 1000 g cal/cm2 through the surface. The results are shown in table 28, which serves to emphasize that the greater part of the energy is absorbed near the surface, particularly in turbid water. If no other processes took place, the temperature between the surface and 1 m would increase in the clearest water by 6.24°, and in the most turbid water by 7.72°. Between 20 and 21 m the corresponding values would be 0.04° and 0.0003°.

The temperature changes recorded in table 28 show no similarity to those actually occurring in the open oceans, where processes of mixing entirely mask the direct effect of absorption, but in some small, landlocked

bodies of water the temperature changes at subsurface depths may be governed mainly by absorption of short-wave radiation. Such processes may be observed in the oyster basins on the west coast of Norway, the temperature characteristics of which were described by Helland-Hansen and studied in detail by Gaarder and Spärck (1932). These basins are in communication with the open sea through narrow and shallow openings, but during winter storms a complete exchange of water often takes place between the basins and the outside. In the ensuing spring, after rains, which cause considerable run-off, the surface layer in the basins will be replaced by fresh or brackish water spreading over the sea water in the deeper portion of the basins and forming a cover that prevents further exchange between the deeper water and the outside sea. Owing to the difference in salinity the density of the deeper water will be much higher than the density of the surface layer. During summer the incoming radiation will be absorbed both in the fresh water on top and in the underlying sea water, and the temperature will rise within both layers. Within the top layer the ordinary convection currents develop, and the temperature is controlled mainly by the air temperature, but owing to the greater salinity of the lower layer the temperature of the deeper water can rise to high values without leading to unstable stratification, and the effect of absorption becomes apparent, because no other processes are of importance.

Interval of depth (m) Oceanic water Coastal water
Clearest Average Average Turbid
0–1 6.24 6.48 7.32 7.72
1–2 0.610 0.720 0.970 0.960
5–6 .236 .282 .164 .120
10–11 .104 .096 .030 .0140
20–21 .040 .030 .0016 .0003
50–51 .0096 .0024 .0534 .0715
100–101 .0016 .0411

Fig. 24 shows the vertical distribution of temperature on June 30 and July 15 in a basin that was examined by Gaarder and Spärck. The days in the period between the stated dates were clear and no rain fell. According to Kimball (1928) the diurnal amount of incoming short-wave radiation was about 740 g cal/cm2/day, or about 11,100 g cal/cm2 for

the entire period. If 6 per cent is subtracted for reflection, the amount entering the water would be about 10,400 g cal/cm2. The temperature curves show that of this amount 1630 g cal/cm2, or 15.5 per cent, was absorbed below a depth of 1 m, and thus 84.5 per cent was absorbed between the surface and 1 m. The latter amount was lost by back radiation, heat conduction to the atmosphere, and evaporation, and at present need not be further considered. It is of interest, however, to point out that the great absorption in the upper meter indicates that the turbidity of the water in the basin was greater than that of ordinary coastal waters (table 27, fig. 23) and approached that of turbid lakes. The heating between 1 and 2 m also indicates water of great turbidity, because, of the total amount of 1630 g cal/cm2 reaching 1 m, 630 g cal/cm2 were absorbed in that layer, the corresponding extinction coefficient being 0.488 (cf. table 27). This result agrees roughly with Gaarder and Spärck's statement that a Secchi disc nearly disappeared at a depth of only 2 m. Conditions in this case were unusually clear cut, but even in such basins the effect of absorption is often obscured by processes of heat conduction.


Vertical distribution of temperature in a Norwegian oyster basin on June 30 and July 15, 1927.

Conduction of Heat Through the Ocean Bottom. It has been estimated that the flow of heat through the bottom of the sea amounts to between 50 and 80 g cal/cm2/year (Helland-Hansen, 1930). This amount represents less than one ten-thousandth part of the radiation received at the surface and can generally be neglected when dealing with the heat budget of the oceans. In a few basins, where the deep water is nearly stagnant and where no conduction of heat takes place from above or from the sides, the amount of heat conducted through the bottom may conceivably play a part in determining the distribution of temperature, but so far no such case is known with certainty (p. 739).

Transformation of Kinetic Energy into Heat. The kinetic energy transmitted to the sea by the stress of the wind on the surface and by part of the tidal energy is dissipated by friction and transformed into heat. The energy transmitted by the wind can be estimated at about one ten-thousandth part of the radiation received at the surface and can be neglected. In shallow coastal waters with strong tidal currents the dissipation of tidal energy is so great, however, that it may become of some local importance. Thus, in the Irish Channel, according to Taylor (1919), the dissipation amounts to about 0.002 g cal/cm2/min, or 1050 g cal/cm2/year. The average depth can be taken as about 50 mn, or 5000 cm, and, if the same water remained in the Irish Channel a full year,

the increase in temperature would be about 0.2°C, on an average. Such an effect, however, has not been established, and, as it can be expected in shallow coastal waters only, it is of no significance to the general heat budget of the oceans. A possible case of heating due to dissipation of tidal energy or to conduction of heat through the ocean bottom has been discussed by Sverdrup (1929).

Heating due to chemical processes can be completely disregarded.

The convection of sensible heat and the condensation of water vapor will be dealt with in the discussion concerning the exchange of heat and water with the atmosphere.

Effective Back Radiation from the Sea Surface. The sea surface emits long-wave heat radiation, radiating nearly like a black body, the energy of the outgoing radiation being proportional to the fourth power of the absolute temperature of the surface. At the same time the sea surface receives long-wave radiation from the atmosphere, mainly from the water vapor. A small part of this incoming long-wave radiation is reflected from the sea surface, but the greater portion is absorbed in a small fraction of a centimeter of water, because the absorption coefficients are enormous at long wave lengths. The effective back radiation from the sea surface is represented by the difference between the “temperature radiation” of the surface and the long-wave radiation from the atmosphere, and this effective radiation depends mainly upon the temperature of the sea surface and the water-vapor content of the atmosphere. According to Ångström (1920), the latter is proportional to the local vapor pressure, which can be computed from the relative humidity if the air temperature is known. Over the oceans the air temperature deviates so little from the sea-surface temperature that the vapor pressure can be obtained with sufficient accuracy from the sea-surface temperature and the relative humidity of the air at a short distance above the surface.


Effective back radiation in gm cal/cm2/min from the sea surface to a clear sky. Represented as a function of sea-surface temperature and relative humidity of the air at a height of a few meters.


Ångström (1920) has published a table that summarizes results of observations of effective radiation against a clear sky from a black body of different temperatures and at different vapor pressures. Fig. 25 has been prepared by means of this table, taking into account the small difference between the radiation of a black body and that of a water surface. The figure shows the effective radiation as a function of sea-surface temperature and of relative humidities between 100 per cent and 70 per cent, but the values that can be read off from the graph may be 10 per cent in error, owing to the scanty information upon which the curves are based. It brings out the interesting fact, however, that, owing to the increased radiation from the atmosphere at higher temperatures (higher vapor pressures), the effective back radiation decreases slowly with increasing temperature. At a temperature of 0°C and a relative humidity of 80 per cent, the effective back radiation is 0.188 g cal/cm2/min, and at a temperature of 25° and the same relative humidity it is 0.167 g cal/cm2/min. At a given temperature the effective radiation decreases with increasing humidity, owing to the increased back radiation from the atmosphere. Thus, at a surface temperature of 15° the effective radiation is about 0.180 g cal/cm2/min at a relative humidity of 70 percent, and about 0.163 g cal/cm2/min at a relative humidity of 100 percent.

The values of the effective back radiation at higher temperatures, as obtained by extrapolation of Ångström's data (fig. 25) are greater than those computed from Brunt#x0027;s empirical formula,

where Q′ is the radiation of a black body having the temperature of the sea surface and e is the vapor pressure of the air in millibars. However, in this formula the numerical values of the coefficients are uncertain and are applicable only within a range of e between 4 and 18 mb.

The diurnal and annual variations of the sea-surface temperatures and of the relative humidity of the air over the oceans are small, and the effective back radiation at a clear sky is therefore nearly independent of the time of the day and of the season of the year, in contrast to the incoming short-wave radiation from the sun and the sky, which is subjected to very large diurnal and seasonal variations.

In the presence of clouds the effective back radiation is cut down because the radiation from the atmosphere is increased. The empirical relation can be written

where Q0 is the back radiation at a clear sky and where C is the cloudiness on the scale 1 to 10. A diurnal or annual variation in the cloudiness will
lead to a corresponding variation in the effective back radiation. On an average, the diurnal variation of cloudiness over the oceans is very small and can be neglected, but the annual variation is in some regions considerable. The above equation is applicable to average conditions only, because the reduction of the effective back radiation due to clouds depends upon the altitude and the density of the clouds. Thus, if the sky is completely covered by cirrus, alto-stratus, or strato-cumulus clouds, the effective radiation is about 0.75Q0, 0.4Q0, and 0.1Q0, respectively.

The annual incoming short-wave radiation from the sun and the sky is greater in all latitudes than the outgoing effective back radiation. According to Mosby (1936) the average annual surplus of incoming radiation between latitudes 0 and 10°N is about 0.170 g cal/cm2/min, and between 60° and 70°N about 0.040 g cal/cm2/min. The surplus of radiation must be given off to the atmosphere, and the exchange of heat and water vapor with the atmosphere is therefore equally as important as the processes of radiation in regulating the ocean temperatures and salinities.

The characteristics of the oceans in respect to radiation are very favorable to man. The water surface reflects only a small fraction of the incoming radiation, and the greater part of the radiation energy is absorbed in the water, distributed by processes of mixing over a layer of considerable thickness, and given off to the atmosphere during periods when the air is colder than the sea surface. The oceans therefore exercise a thermostatic control on climate. Conditions are completely changed, however, if the temperature of the sea surface decreases to the freezing point so that further loss of heat from the sea leads to formation of ice, because, when water passes this critical temperature, its thermostatic characteristics are altered in a very unfavorable direction. Sea ice, which soon attains a gray-white appearance owing to enclosed air bubbles, reflects 50 per cent or more of the incoming radiation, and if it is covered by rime or snow the reflection loss increases to 65 per cent, or even to 80 per cent from fresh, dry snow. The snow surface, on the other hand, radiates nearly like a black body, and consequently the heat budget related to processes of radiation, instead of rendering a surplus as it does over the open ocean, shows a deficit until the temperature of the ice surface has been lowered so much that the decreased loss by effective back radiation balances the small fraction of the incoming radiation that is absorbed. The immediate result of freezing is therefore a general lowering of the surface temperature of the ice and a rapid increase of the thickness of the ice. The air that comes in contact with the ice is cooled, and, as this cold air spreads, more ice is formed. Thus, a small lowering of the temperature of the water in high latitudes followed by freezing may lead to a rapid drop of the air temperature and

a rapid increase of the ice-covered area. On the other hand, a small increase of the temperature of air flowing in over an ice-covered sea may lead to melting of the ice at the outskirts and, once started, the melting may progress rapidly. In agreement with this reasoning it has been found that the extent of ice-covered areas in the Barents Sea is a sensitive indicator of small changes in the atmospheric circulation and in the amount of warm water carried into the region by currents (p. 662). It has also been computed that, if the average air temperature in middle and higher latitudes were raised a few degrees, the Polar Sea would soon become an ice-free ocean.

Exchange of Heat between the Atmosphere and the Sea. The amount of heat that in unit time is carried away from the sea surface through a unit area is equal to

where cp is the specific heat of the air, A is the eddy conductivity, – dϑ/dz is the temperature gradient of the air (the lapse rate), which is positive when the temperature decreases with height, and γ is the adiabatic lapse rate. Very near the sea surface, γ can be neglected as small compared to dϑ/dz. The term cpA enters instead of the coefficient of heat conductivity of the air as determined in the laboratory because the air is nearly always in turbulent motion (p. 92). The state of turbulence varies, however, with the distance from the sea surface, because at the surface itself the eddy motion must be greatly reduced. As a consequence, under steady conditions, when the same amount of heat passes upward through every cross section of a vertical column, the temperature changes rapidly with height near the sea surface and more slowly at greater distance. The product –cpAdϑa/dz remains constant, and, as cpA increases rapidly with height, –dϑa/dz must decrease.

Detailed and accurate temperature measurements in the lowest meters of air over the ocean have not yet been made, because the hull and masts of a vessel disturb the normal distribution of temperature to such an extent that values observed at different levels on board a vessel are not representative of the undisturbed conditions. The few measurements that have been attempted indicate, however, that the general distribution as outlined above is encountered.

The sea surface must be warmer than the air at a small distance above the surface if heat is to be conducted from the sea to the air. When such conditions prevail, the air is heated from below, the stratification of the air becomes unstable, and the turbulence of the air becomes intense (p. 92). If the sea surface is very much warmer than the air, as may be

the case when cold continental air flows out over the sea in winter, the heating from below may be so intense that rapid convection currents develop, leading to such violent atmospheric disturbances as thunderstorms. It is not intended here to enter upon the meteorological aspects of the heat exchange, but the point which is emphasized is that an appreciable conduction of heat from the sea to the atmosphere takes place only when the sea surface is warmer than the air. One might assume that, vice versa, an appreciable amount of heat would be conducted to the sea surface when warmer air flows over a cold sea, but this is not the case, because under such conditions the air is cooled from below, the stratification of the air becomes stable, and the turbulence (and consequently the eddy conductivity of the air) is greatly reduced.

It has been found (p. 128) that on an average the sea surface is slightly warmer than the overlying air and therefore loses heat by conduction. So far, no detailed studies have been made, but Ångström has estimated that only about 10 per cent of the total heat surplus is given off to the atmosphere by conduction and that 90 per cent is used for evaporation. Other estimates indicate that these figures are approximately correct (p. 117). Thus, evaporation is of much greater importance to the heat balance of the oceans than is the transfer of sensible heat. Evaporation will therefore be dealt with in greater detail.

Evaporation from the Sea

The Process of Evaporation. The vapor tension at a flat surface of pure water depends on the temperature of the water. The salinity decreases the tension slightly, the empirical relation between vapor tension and salinity being (p. 66)

where es is the vapor tension over sea water, ed is the vapor tension over distilled water of the same temperature, and S is the salinity in parts per mille. In the open ocean the relation is approximately es = 0.98 ed. Table 29 contains the vapor tension in millibars over water of salinity 35.00 ‰ and at the stated temperatures.

Air in which the vapor tension is less than that over water of the same temperature is undersaturated with moisture, and air in which the vapor tension exactly equals that over a water surface of the same temperature is saturated with moisture. In absolutely pure air the vapor pressure can be above the saturation value, but generally the air contains “nuclei” on which the vapor is condensed when the vapor tension reaches the value corresponding to that over water of the same temperature. Under these conditions the vapor tension in the air cannot be further increased, and in meteorology one therefore uses the

term “maximum vapor tension” at a given temperature. The maximum vapor tension at which condensation takes place can be reached either by adding water vapor to air of a given temperature or by decreasing the temperature of air of a given moisture content. In the latter case condensation begins at the temperature that is called the “dew point.”

Temperature (°C) Vapor pressure (mb) Temperature (°C) Vapor pressure (mb)
−2 5.19 18 20.26
−1 5.57 19 21.57
0 5.99 20 22.96
1 6.44 21 24.42
2 6.92 22 25.96
3 7.43 23 27.59
4 7.98 24 29.30
5 8.56 25 31.12
6 9.17 26 33.01
7 9.83 27 35.02
8 10.52 28 37.13
9 11.26 29 39.33
10 12.05 30 41.68
11 12.88 31 44.13
12 13.76 32 46.71
13 14.70
14 15.69
15 16.74
16 17.85
17 19.02

In discussing the process of evaporation it is more rational to consider not the vapor pressure but the specific humidity, f—that is, the mass of water vapor per unit mass of air. The amount of water vapor, F, which per second is transported upward through a surface of cross section 1 cm2 is, then, – Adf/dz, where A is the eddy conductivity and – df/dz is the vertical gradient of the specific humidity, which is positive when the specific humidity decreases with height. If the vapor pressure, e, is introduced, the equation becomes approximately


where p is the atmospheric pressure. The heat needed for evaporation at the surface is
where Lϑ is the heat of vaporization at the temperature of the surface, ϑ (p. 62).

The ratio between the amounts of heat given off to the atmosphere as sensible heat (p. 114) and used for evaporation is

The last expression is obtained by introducing cp = 0.240 and L = 585. Thus, the ratio R depends mainly upon the ratio between the temperature and humidity gradients in the air at a short distance from the sea surface. These gradients are difficult to measure but can be replaced approximately by the difference in temperature and vapor pressure at the sea surface and the corresponding values in the air at a height of a few meters:
This ratio was derived in a different manner by Bowen (1926), and is often referred to as the “Bowen ratio.”

Values of the ratio R can be computed from climatological charts of the oceans, but a comprehensive study has not been made. Calculations based on data contained in the Atlas of Climatic Charts of the Oceans, published by the U. S. Weather Bureau (1938), show that the ratio varies from one part of the ocean to the other. As a rule, the ratio is small in low latitudes, where it remains nearly constant throughout the year, but is greater in middle latitudes, where it reaches values up to 0.5 in winter and in some areas drops to –0.2 in summer. A negative value indicates that heat is conducted from the atmosphere to the sea. On an average, the value for all oceans appears to lie at about 0.1, meaning that about 10 per cent of the heat surplus that the oceans receive by radiation processes is given off as sensible heat, whereas about 90 per cent is used for evaporation (p. 115).

There are certain points regarding the character of the evaporation which need to be emphasized. If the water is warmer than the air, the vapor pressure at the sea surface remains greater than that in the air, and evaporation can always take place and will be greatly facilitated in these circumstances because the turbulence of the air will be fully developed owing to the unstable stratification of the very lowest layers (p. 92). It must therefore be expected that the greatest evaporation occurs when cold air flows over warm water. If the air is much colder

than the water, the air may become saturated with water vapor, and fog or mist may form over the water surfaces. Such fog is common in the fall over ponds and small lakes during calm, clear nights. When a wind blows, the moisture will be carried upward, but streaks and columns of fog are often visible over lakes or rivers and are commonly described as “smoke.” The process can occasionally be observed near the coast, but not over the open ocean, because the necessary great temperature differences are rapidly eliminated as the distance from the coast increases.


Left: The difference, air minus sea-surface temperature, and the prevailing wind-direction over the Grand Banks of Newfoundland in March, April, and May. Right: Percentage frequency of fog in the same months.

When the sea surface is colder than the air, evaporation can take place only if the air is not saturated with water vapor. In this case turbulence is reduced and evaporation must stop when the vapor content of the lowest layer of the atmosphere has reached such a value that the vapor pressure equals that at the sea surface. If warm, moist air passes over a colder sea surface, the direction of transport will be reversed and condensation will take place on the sea surface in such a way that heat is brought to the surface and not carried away from it. Owing to the fact that this process can take place only when the air is warmer than the sea and that then turbulence is greatly reduced, one can expect that condensation of water vapor on the sea will not be of great importance, but it should be borne in mind that this process can and does take place when conditions are right. In these circumstances, contact with the sea and conduction lower the air temperature to the dew point for a considerable distance above the sea surface. Condensation takes place in the air and fog is formed, “advection” fog that is commonly encountered over the sea. The relation between the frequency of fog or mist and the difference between sea-surface and air temperatures are well illustrated by charts in the Atlas of Climatic Charts of the Oceans (1938). As an example, fig. 26 shows the frequency of fog, the difference between air and sea-surface temperature, and the prevailing wind direction over the Grand Banks of Newfoundland in March, April, and May. It can be concluded that in spring, when the water is colder than the air, no

evaporation takes place in this region, but in the fall and winter, when the water is warmer, evaporation must be great.

In middle and higher latitudes the sea surface in winter is mostly warmer than the air, and hence one must expect the evaporation then to be at its maximum. This conclusion appears contrary to common experience that evaporation from heated water is greater than that from cold water, but the contradiction is only apparent, because greatest evaporation always occurs when a water surface is warmer than the air above it, which is exactly what happens in winter.

Observations and Computations of Evaporation. Present knowledge of the amount of evaporation from the different parts of the oceans is derived partly from observations and partly from computations based on consideration of the heat balance.

Observations have been made by means of pans on board ship, but such observations give values of the evaporation from the sea surface that are too high, partly because the wind velocity is higher at the level of the pan than at the sea surface, and partly because the difference between vapor pressure in the air and that of the evaporating surface is greater at the pan than at the sea surface. Analyzing the decrease of the wind velocity and the increase of the vapor pressure between the average level of pans used on shipboard and a level a few centimeters above the sea surface, Wüst (1936) arrived at the conclusion that the measured values had to be multiplied by 0.53 in order to represent the evaporation from the sea surface.

In computing the evaporation on the basis of the heat balance, one has to begin with the equation (p. 101)

Introducing the ratio R = Qh/Qe, and converting the evaporation, E, into centimeters by dividing Qe by the latent heat of vaporization, L, one obtains
In this form the equation representing the heat balance has found wide application for computation of evaporation. The result gives the evaporation in centimeters during the time intervals to which the values Qs, and so on, apply, provided these are expressed in gram calories.

A second method for computing the evaporation from the oceans has been suggested by Sverdrup (1937), who, on the basis of results in fluid mechanics as to the turbulence of the air over a rough surface, established a formula for the evaporation, using in part constants that had been determined by laboratory experiments and in part constants that were obtained from the character of the variation of vapor pressure with

increasing height above the sea surface. Similar but more complicated formulae have been derived by Millar (1937) and by Montgomery (1940).

The exact formulae are not well suited for numerical computation, but at wind velocities between 4 and 12 m/sec, the mean annual evaporation in centimeters can be found approximately from the simple relation

where ēw represents the average vapor pressure in millibars at the sea surface as derived from the temperature and salinity of the sea, ēa represents the average vapor pressure in the air at a height of 6 m above the sea surface, and ū is the average wind velocity in meters per second at the same height.

Average Annual Evaporation from the Oceans. On the basis of pan measurements conducted in different parts of the ocean, Wüst (1936) found that the average evaporation from all oceans amounts to 93 cm/year, and he considers this value correct within 10 to 15 per cent. W. Schmidt (1915) computed the evaporation by means of the preceding equation for E, in which the terms Qϑ and Qv can be omitted in considering the oceans as a whole. Schmidt introduced high values of R, and on the basis of the available data as to incoming radiation and back radiation he found a total evaporation of 76 cm/year. A revision based on more recent measurements of radiation (Mosby, 1936) and use of R = 0.1 resulted in a value of 106 cm/year. The latter value represents an upper limit, and may be 10 to 15 per cent too high, whence it appears that Wüst's result is nearly correct.

It is of interest in this connection to give some figures regarding the relation between evaporation and precipitation over the oceans, the land areas, and the whole earth (according to Wüst, 1936). The total evaporation from the oceans amounts to 334,000 km3/year, of which 297,000 km3 returns to the sea in the form of precipitation, and the difference, 37,000 km3, must be supplied by run-off, since the salinity of the oceans remains unchanged. The total amount of precipitation falling on the land is 99,000 km3, of which amount a little over one third, 37,000 km3, is supplied by evaporation from the oceans and 62,000 km3 is supplied by evaporation from inland water areas or directly from the moist soil. For the sake of comparison it may be mentioned that the capacity of Lake Mead, above Boulder Dam, is 45 km3.

Evaporation in Different Latitudes and Longitudes. From pan observations at sea, Wüst has derived average values of the evaporation from the different oceans in different latitudes (table 30, p. 123). By means of the energy equation one can compute similar annual values, assuming that the net transport of heat by ocean currents can be neglected. Such a computation has been carried out for the Atlantic

Ocean, making use of Kimball#x0027;s data (1928) as to the incoming radiation and the observed temperatures and humidities for determining effective back radiation. In fig. 27 are shown Wüst#x0027;s values of the annual evaporation between latitudes 50°N and 50°S in the Atlantic Ocean and the corresponding values as derived from the energy equation. The agreement is very satisfactory. The low evaporation in the equatorial regions that both curves show can be ascribed to the higher relative humidities and the lower wind velocities of that area, if one considers the processes of evaporation, or it can be ascribed to the effect of the prevailing cloudiness if one considers the energy relations. The great evaporation in the areas of subtropical anticyclones appears clearly, but in the Southern Hemisphere the observations give the highest values of the evaporation nearer to the Equator than do the computations. The discrepancy may be due to the fact that in the course of a year the subtropical anticyclone changes its distance from the Equator and that the observations have not been distributed evenly over the year. The energy equation has also been used by McEwen (1938) for computing values of evaporation over the eastern Pacific Ocean between latitudes 20°N and 50°N. His figures agree with those obtained by Wüst for the same latitudes.


Annual evaporation from the Atlantic Ocean between latitudes 50°N and 50°S. Thin curve based on observations (Wüst, 1936) and heavy curve on computations, using the energy equation.

It appears that the average annual values of the evaporation in different latitudes are well established, but the evaporation also varies from the eastern to the western parts of the oceans and with the seasons. These variations are of great importance to the circulation of the atmosphere, because the supply of water vapor that later on condenses and gives off its latent heat represents a large portion of the supply of energy. So far, none of the details are known, but it is possible that approximate values of the evaporation from different parts of the ocean and in different seasons can be found by means of the method proposed by Sverdrup (1937) and used by Jacobs (1942).


Annual Variation of Evaporation. The character of the annual variation of evaporation can be examined by means of the energy equation (Sverdrup, 1940):


The quantity Qϑ can be computed if the annual variation of temperature due to processes of heating and cooling is known at all depths where such annual variations occur. The annual variation of temperature at the surface has been examined, but only few data are available from subsurface depths, the most reliable being those which have been compiled by Helland-Hansen (1930) from an area in the eastern North Atlantic with its center in 47°N and 12°W (p. 132). The radiation income in that area can be obtained from Kimball#x0027;s data (1928), the back radiation can be found by means of the diagram in fig. 25, and the transport by currents, Qv, can be neglected. In fig. 28A are represented the annual variation of the net surplus of radiation, Qr, the annual variation of the amount of heat used for changing the temperature of the water, Qϑ, and the difference between these two amounts, Qa, which represents the total amount of heat given off to the atmosphere. The greater part of the last amount is used for evaporation, and the curve marked Qa represents, therefore, approximately the annual variation of the evaporation, which shows a maximum in the fall and early winter, a secondary minimum in February, followed by a secondary maximum in March, and a low minimum in summer. In June and July no evaporation takes place. The total evaporation during the year is about 80 cm.


(A) Annual variation in the total amount of heat, qa, given off to the atmosphere in an area of the North Atlantic (about 47°N, 12°W). (B) Corresponding diurnal variation near the Equator in the Atlantic Ocean. (For explanation of symbols, see text.)

This example illustrates a method of approach that may be applied, but so far the necessary data for a more complete study are lacking. The result that the evaporation is at a minimum in summer and at a maximum in fall and early winter is in agreement with the conclusions that were drawn when discussing the process of evaporation in general.

Latitude Atlantic Ocean Indian Ocean Pacific Ocean All Oceans
S (‰) E (cm/yr) P (cm/yr) EP (cm/yr) S (‰) E (cm/yr) P (cm/yr) EP (cm/yr) S (‰) E (cm/yr) P (cm/yr) EP (cm/yr) S (‰) E (cm/yr) P (cm/yr) EP (cm/yr)
40°N 35.80 94 76 18 33.64 94 93 1 34.54 94 93 1
35 36.46 107 64 43 34.10 106 79 27 35.05 106 79 27
30 36.79 121 54 67 34.77 116 65 51 35.56 120 65 55
25 36.87 140 42 98 35.00 127 55 72 35.79 129 55 74
20 36.47 149 40 110 (35.05) (125) (74) (51) 34.88 130 62 68 35.44 133 65 68
15 35.92 145 62 83 (35.07) (125) (73) (52) 34.67 128 82 46 35.09 130 82 48
10 35.62 132 101 31 (34.92) (125) (88) (37) 34.29 123 127 − 4 34.72 129 127 2
5 34.98 105 144 − 39 (34.82) (125) (107) (18) 34.29 102 (177) (-75) 34.54 110 177 − 67
0 35.67 116 96 20 35.14 125 131 − 6 34.85 116 98 18 35.08 119 102 17
5°S 35.77 141 42 99 34.93 121 167 − 46 35.11 131 91 40 35.20 124 91 33
10 36.45 143 22 121 34.57 99 156 − 57 35.38 131 96 35 35.34 130 96 34
15 36.79 138 19 119 34.75 121 83 38 35.57 125 85 40 35.54 134 85 49
20 36.54 132 30 102 35.15 143 59 84 35.70 121 70 51 35.69 134 70 64
25 36.20 124 40 84 35.45 145 46 99 35.62 116 61 55 35.69 124 62 62
30 35.72 116 45 71 35.89 134 58 76 35.40 110 64 46 35.62 111 64 47
35 35.35 99 55 44 35.60 121 60 61 35.00 97 64 33 35.32 99 64 35
40 34.65 81 72 9 35.10 83 73 10 34.61 81 84 − 3 34.79 81 84 − 3
45 34.19 64 73 − 9 34.25 64 79 − 15 34.32 64 85 − 21 34.14 64 85 − 21
50 33.94 43 72 − 29 33.87 43 79 − 36 34.16 43 84 − 41 33.99 43 84 − 41


Diurnal Variation of Evaporation. The diurnal variation of evaporation can be examined in a similar manner, but at the present time suitable data are available only at four Meteor stations near the Equator in the Atlantic Ocean (Defant, 1932; Kuhlbrodt and Reger, 1938). In fig. 28B the curves marked Qr and Qϑ correspond to the similar curves in fig. 28A, and the difference between these, Qa, shows the amount of heat lost during twenty-four hours, which is approximately proportional to the evaporation. The diurnal variation of evaporation in the Tropics appears to have considerable similarity to the annual variation in middle latitudes, and is characterized by a double period with maxima in the late forenoon and the early part of the night and minima at sunrise and in the early afternoon hours. It is possible that the afternoon minimum appears exaggerated, owing to uncertainties as to the absolute values of Qr and Qϑ. The total diurnal evaporation was 0.5 cm, but the sky was nearly clear on the four days that were examined and the average diurnal value is therefore smaller. The double diurnal period of evaporation appears to be characteristic of the Tropics, but in middle latitudes a single period with maximum values during the night probably dominates.

Salinity and Temperature of the Surface Layer

The Surface Salinity. In all oceans the surface salinity varies with latitude in a similar manner. It is at a minimum near the Equator, teaches a maximum in about latitudes 20°N and 20°S, and again decreases toward high latitudes.


(A) Average values for all oceans of surface salinity and the difference, evaporation minus precipitation, plotted against latitude. (B) Corresponding values of surface salinity and the difference, evaporation minus precipitation, plotted against each other (according to Wüst, 1936).

Table 30 contains average values of the surface salinity, the evaporation, the precipitation, and the difference between the last quantities for the three major oceans and for all combined, according to Wüst (1936).

On the basis of these values, Wüst has shown that for each ocean the deviation of the surface salinity from a standard value is directly proportional to the difference between evaporation, E, and precipitation, P. In fig. 29A are plotted the surface salinities for all oceans and the difference, EP, in centimeters per year, as functions of latitude; the corresponding values of salinity and the difference, EP, are plotted against each other. If the values of 5°N are omitted, because they disagree with all others, the values fall nearly on a straight fine leading to the empirical relationship

Wüst points out that such an empirical relationship is found because the surface salinity is mainly determined by three processes: decrease of salinity by precipitation, increase of salinity by evaporation, and change of salinity by processes of mixing. If the surface waters are mixed with water of a constant salinity, and if this constant salinity is represented by S0, the change of salinity due to mixing must be proportional to S0 – S, where S is the surface salinity. The change of salinity due to processes of evaporation and precipitation must be proportional to the difference (EP). The local change of the surface salinity must be zero; that is,


As this simple formula has been established empirically, it must be concluded that the surface water is generally mixed with water of a salinity which, on an average, is 34.6 ‰. This value represents approximately the average value of the salinity at a depth of 400 to 600 m, and it appears therefore that vertical mixing is of great importance to the general distribution of surface salinity. This concept is confirmed by the fact that the standard value of the salinity differs for the different oceans. For the North Atlantic and the North Pacific, Wüst obtains similar relationships, but the constant term, S0, has the value 35.30 ‰ in the North Atlantic and 33.70 ‰ in the North Pacific Ocean. The corresponding average values of the salinity at a depth of 600 m are 35.5 ‰ and 34.0 ‰, respectively. For the South Atlantic and the South Pacific Oceans, Wüst finds S0 = 34.50 ‰ and 34.64 ‰, respectively, and the average salinity at 600 m in both oceans is about 34.5 ‰. In these considerations the effect of ocean currents on the distribution of surface salinity has been neglected, and the simple relations obtained indicate that transport by ocean currents is of minor importance as far as average conditions are concerned. The difference between evaporation and precipitation, EP, on the other hand, is of primary importance, and,

because this difference is closely related to the circulation of the atmosphere, one is led to the conclusion that the average values of the surface salinity are to a great extent controlled by the character of the atmospheric circulation.

The distribution of surface salinity of the different oceans is shown in chart VI, in which the general features that have been discussed are recognized, but the details are so closely related to the manner in which the water masses are formed and to the types of currents that they cannot be dealt with here.

Periodic Variations of the Surface Salinity. Over a large area, variations in surface salinity depend mainly upon variations in the difference between evaporation and precipitation. From Böhnecke#x0027;s monthly charts (1938) of the surface salinity of the North Atlantic Ocean, mean monthly values have been computed for an area extending between latitudes 18° and 42°N, omitting the coastal areas in order to avoid complications due to shifts of coastal currents. The results of this computation (fig. 30) show the highest average surface salinity, 36.70‰ in March, and the lowest, 36.59 ‰, in November. The variations from one month to another are irregular, but on the whole the salinity is somewhat higher in spring than it is in the fall.


Annual variation of surface salinity in the North Atlantic Ocean between latitudes 18°N and 42°N.

Harmonic analysis leads to the result

and thus
Because ∂S/∂t is proportional to EP, it follows that the excess of evaporation over precipitation is at a minimum at the end of June and at a maximum at the end of December. This annual variation corresponds closely to the annual variation of evaporation (p. 122), for which reason it appears that in the area under consideration the annual variation of the surface salinity is mainly controlled by the variation in evaporation during the year. For a more exact examination, subsurface data are needed, but nothing is known as to the annual variation of salinity at subsurface depths.

More complicated conditions are encountered in the northwestern part of the Atlantic Ocean, where, according to G. Neumann (1940),

between the Azores and Newfoundland the annual variation of salinity has the character of a disturbance that originates to the southwest of Newfoundland and spreads toward the east and east-southeast. Off Newfoundland, the amplitude is 0.37 ‰ and the maximum occurs about March 1 (phase angle equals −60 degrees). Toward the east and eastsoutheast the amplitude decreases and the maximum is reached later and later, as if the disturbance progressed like a wave that was subject to damping. On the assumption that this damping is caused by horizontal mixing, Neumann finds that the corresponding coefficient of mixing lies between 2 × 108 and 5 × 108 cm2/sec.

From the open ocean no data are available as to the diurnal variation of the salinity of the surface waters, but it may be safely assumed that such a variation is small, because neither the precipitation nor the evaporation can be expected to show any considerable diurnal variation.

North latitude Atlantic Ocean Indian Ocean Pacific Ocean South latitude Atlantic Ocean Indian Ocean Pacific Ocean
70°–60° 5.60 ...... ...... 70°–60° − 1.30 − 1.50 − 1.30
60–50 8.66 5.74 60–50 1.76 1.63 5.00
50–40 13.16 9.99 50–40 8.68 8.67 11.16
40–30 20.40 ...... 18.62 40–30 16.90 17.00 16.98
30–20 24.16 26.14 23.38 30–20 21.20 22.53 21.53
20–10 25.81 27.23 26.42 20–10 23.16 25.85 25.11
10–0 26.66 27.88 27.20 10–0 25.18 27.41 26.01

Surface Temperature. The genera1 distribution of surface temperature cannot be treated in a manner similar to that employed by Wüst when dealing with the salinity, because the factors controlling the surface temperature are far more complicated. The discussion must be confined to presentation of empirical data and a few general remarks.

Table 31 contains the average temperatures of the oceans in different latitudes according to Krümmel (1907), except in the case of the Atlantic Ocean, for which new data have been compiled by Böhnecke (1938). In all oceans the highest values of the surface temperature are found somewhat to the north of the Equator, and this feature is probably related to the character of the atmospheric circulation in the two hemispheres. The region of the highest temperature, the thermal Equator, shifts with the season, but in only a few areas is it displaced to the Southern Hemisphere at any season. The larger displacements (Schott, 1935, and Böhnecke, 1938) are all in regions in which the surface currents change

during the year because of changes in the prevailing winds, and this feature also is therefore closely associated with the character of the atmospheric circulation. The surface temperatures in the Southern Hemisphere are generally somewhat lower than those in the Northern, and again the difference can be ascribed, to difference in the character of the prevailing winds, and perhaps also to a widespread effect of the cold, glacier-covered Antarctic Continent.

The average distribution of the surface temperature of the oceans in February and August is shown in charts II and III. Again the distribution is so closely related to the formation of the different water masses and the character of the currents that a discussion of the details must be postponed.

Difference Between Air and Sea-Surface Temperatures. It was pointed out that in all latitudes the ice-free oceans received a surplus of radiation, and that therefore in all latitudes heat is given off from the ocean to the atmosphere in the form of sensible heat or latent heat of water vapor. The sea-surface temperature must therefore, on an average, be higher than the air temperature. Observations at sea have shown that such is the case, and from careful determinations of air temperatures over the oceans it has furthermore been concluded that the difference, air minus sea-surface temperature, is greater than that derived from routine ships' observations. In order to obtain an exact value, it is necessary to measure the air temperature on the windward side of the vessel at a locality where no eddies prevail, but where the air reaches the thermometer without having passed over any part of the vessel. For measurements of the temperature a ventilated thermometer must be used. According to the Meteor observations (Kuhlbrodt and Reger, 1938) the air temperature over the South Atlantic Ocean between latitudes 55°S and 20°N is on an average 0.8 degree lower than that of the surface, whereas in the same region the atlas of oceanographic and meteorological observations published by the Netherlands Meteorological Institute gives an average difference of only 0.1 degree. The reason for this discrepancy is that the air temperatures as determined on commercial vessels are on an average about 0.7 degree too high because of the ships' heat. The result as to the average value of the difference, υw – υa, is in good agreement with results obtained on other expeditions when special precautions were taken for obtaining correct air temperatures. Present atlases of air and sea-surface temperatures have been prepared from the directly observed values on board commercial vessels without application of a correction to the air temperatures. This correction is so small that it is of minor importance when the atlases are used for climatological studies, but in any studies that require knowledge of the exact difference between air- and sea-surface temperatures it is necessary to be aware of the systematic error in the air temperature.


The difference of 0.8 degree between air and surface temperatures, as derived from the Meteor observations, is based on measurements of air temperature at a height of 8 m above sea level. At the very sea surface the air temperature must coincide with that of the water, and consequently the air temperature decreases within the layers directly above the sea. The most rapid decrease takes place, however, very close to the sea surface, and at distances greater than a few meters the decrease is so slow that it is immaterial whether the temperature is measured at 6, 8, or 10 m above the surface. The height at which the air temperature has been observed on board a ship exercises a minor influence, therefore, upon the accuracy of the result, and discrepancies due to differences in the height of observations are negligible compared to the errors due to inadequate exposure of the thermometer.

The statement that the air temperature is lower than the water temperature is correct only when dealing with average conditions. In any locality the difference, υw – υa, generally varies during the year in such a manner that in winter the air temperature is much lower than the sea-surface temperature, whereas in summer the difference is reduced or the sign is reversed. The difference also varies from one region to another according to the character of the circulation of the atmosphere and of the ocean currents. These variations are of great importance to the local heat budget of the sea because the exchange of heat and vapor between the atmosphere and the ocean depends greatly upon the temperature difference.

It was shown that the amount of heat given off from the ocean to the surface is, in general, great in winter and probably negligible in summer. Owing to this annual variation in the heat exchange, one must expect that in winter the air over the oceans is much warmer than the air over the continents but in summer the reverse conditions should be expected. That such is true is evident from a computation of the average temperature of the air between latitudes 20°N and 80°N along the meridian of 120°E, which runs entirely over land, and along the meridian of 20°W, which runs entirely over the ocean (von Hann, 1915, p. 146). In January the average temperature along the “land meridian” of 120°E is −15.9 degrees C, but along the “water meridian” of 20°W it is 6.3 degrees. In July the corresponding values are 19.4 and 14.6 degrees, respectively. Thus, in January the air temperature between 20°N and 80°N over the water meridian is 22.2 degrees higher than that over the land meridian, whereas in July it is 4.8 degrees lower. The mean annual temperature is 7.0 degrees higher along the water meridian.

Annual Variation of Surface Temperature. The annual variation of surface temperature in any region depends upon a number of factors, foremost among which are the variation during the year of the radiation income, the character of the ocean currents, and of the prevailing

winds. The character of the annual variation of the surface temperature changes from one locality to another, but a few of the general features can be pointed out. The heavy curves in fig. 31 show the average annual range of the surface temperature in different latitudes in the Atlantic, the Indian, and the Pacific Oceans. The range represents the difference between the average temperatures in February and August and is derived for the Atlantic from Böhnecke's tables (1938), and, for the Indian and the Pacific Oceans, from the charts published by Schott (1935). Thin lines in the same figure show the range of the radiation income as derived from Kimball's maps (1928). The curves bring out two characteristic features. In the first place they show that the annual range of the surface temperature is much greater in the North Atlantic and the North Pacific Oceans than in the southern oceans. In the second place they show that in the southern oceans the temperature range is definitely related to the range in radiation income, whereas in the northern oceans no such definite relation appears to exist. The great ranges in the northern oceans are associated with the character of the prevailing winds and, particularly, with the fact that cold winds blow from the continents toward the ocean and greatly reduce the winter temperatures. Near the Equator a semiannual variation is present, corresponding to the semiannual period of radiation income, but in middle and higher latitudes the annual period dominates.


Average annual ranges of surface temperature in the different oceans plotted against latitude (heavy curves) and corresponding ranges in the radiation income (thin curves).


Annual Variation of Temperature in the Surface Layers. At subsurface depths the variation of temperature depends upon four factors: (1) variation of the amount of heat that is directly absorbed at different depths, (2) the effect of heat conduction, (3) variation in the currents related to lateral displacement of water masses, and (4) the effect of vertical motion. The annual variation of temperature at subsurface depths cannot be dealt with in a general manner, owing to lack of data, but it is again possible to point out some outstanding characteristics, using two examples from the Pacific and one from the Atlantic Ocean. The effects of all four of the important factors are illustrated in fig. 32A, which shows the annual variation of temperature at the surface and at depths of 25, 50, and 100 m at Monterey Bay, California (Skogsberg, 1936). Skogsberg divides the year into three periods: the period of the Davidson Current, lasting from the middle of November to the middle of February; the period of upwelling, between the middle of February and the end of July; and the oceanic period, from the end of July to the middle of November. The California Current off Monterey Bay during the greater part of the year is directed to the south, but during winter, from the middle of November to the middle of February, an inshore flow to the north, the Davidson Current, is present (p. 724). The water of this inshore flow is characterized by relatively high and uniform temperature and appears in the annual variation of temperature as warm water at subsurface depths. The upper homogeneous layer is relatively thick, as is evident from the fact that the temperature is nearly the same at 25 m as it is at the surface, and that at 50 m it is only slightly lower. At the end of February the California Current again reaches to the coast and, under the influence of the prevailing northwesterly winds, an overturn of the upper layers

takes place that is generally described as upwelling (p. 501). During the period of upwelling, vertical motion near the coast brings water of relatively low temperature toward the surface. Consequently, the temperatures at given depths decrease when the upwelling begins. This decrease is brought out in fig. 32A by the downward trend of the temperature at 25, 50, and 100 m, at which depths the minimum temperature is reached at the end of May. The much higher temperature at the surface as compared to that at 25 m shows that a thin surface layer is subject to heating by radiation, and from the variation of temperature at 10 m, which is shown by a thin line, it is evident that the effect of heating is limited to the upper 10 m. As the upwelling gradually ceases toward the end of August, a sharp rise in temperature takes place both at the surface and at subsurface depths, and the peaks shown by the temperature curves in September are results of heating and conduction and intrustion of offshore water. Thus, the annual march of temperature can be explained from changes in currents, vertical motion associated with upwelling, seasonal heating and cooling, and heat conduction.


(A) Annual variation of temperature at different depths in Monterey Bay, California. (B) Annual variation of temperature at different depths in the Kuroshio off the South Coast of Japan.

The annual variation of temperature in the Kuroshio off the south coast of Japan (Koenuma, 1939), as shown in fig. 32B, gives an entirely different picture. The annual variation has the same character at all depths between the surface and 100 m, with a minimum in late winter and a maximum in late summer or early fall, but the range of the variation decreases with depth, and the time of maximum temperature occurs later with increasing depth. From the course of the curves it may be concluded that the annual variation is due to heating and cooling near the surface and is transmitted to greater depths by processes of conduction (p. 136). This appears to be correct, but the heating and cooling is only partly caused by variations in the net radiation, and it also depends on excessive cooling in winter by cold and dry winds blowing toward the sea (Sverdrup, 1940).

In order to be certain that observed temperature variations are related to processes of heating and cooling only, it is necessary to examine whether the water in a given locality is of the same character throughout the year. For this purpose Helland-Hansen (1930) developed a method that is applicable in areas in which it is possible to establish a definite relation between temperature and salinity (p. 142). He assumed that any temperature value above or below that determined by the temperature-salinity relation can be considered as resulting from heating or cooling of the water, and he used the method within three areas in the eastern North Atlantic. Fig. 33 shows the curves which he determined for an area off the Bay of Biscay with its center in approximately 47°N and 12°W. The character of the curves, the reduction of the range, and the displacement of the times of maxima clearly show that one has to deal with heat conduction. In this case the variation in the heat content

corresponds nearly to the variation in net radiation, whereas in the Kuroshio the additional effect of excessive cooling in winter by winds from the continent leads to much greater variations in temperature and heat content.

These examples serve to illustrate different types of annual variation of temperature that may be encountered in different localities and also to stress the fact that conclusions as to the temperature variations associated with processes of local heating are valid only if the data are such that the influences of shifting currents and vertical motion can be eliminated.

Diurnal Variation of Surface Temperature. The range of the diurnal variation of surface temperature of the sea is not more, on an average, than 0.2 to 0.3 degree. Earlier observations gave somewhat higher values, particularly in the Tropics, but new, careful measurements and reexamination of earlier data in which doubtful observations have been eliminated have shown that the range of diurnal variation is quite small. Meinardus (Kuhlbrodt and Reger, 1938, p. 301–302) summarizes his examination of a large number of data by stating, “in general, the diurnal variation of water temperature in lower latitudes can be represented by a sine curve with extreme values between 2:30h and 3h and between 14:30h and 15h, and a range of 0.3° to 0.4°. In higher latitudes the extreme values come later and the range is even smaller.” The Meteor observations give ranges of only 0.2 to 0.3 degree in the Tropics. The Meteor data and the Challenger data, which have been discussed by Wegemann (1920), both show that close to the Equator the diurnal variation of surface temperature is somewhat unsymmetrical, the temperature increasing rapidly after sunrise and decreasing slowly after sunset, but at greater distances from the Equator the curve becomes somewhat more symmetrical.


Annual variation of temperature at different depth off the Bay of Biscay in approximately 47°N and 12°W.


The changes through the year of the range of diurnal variation of surface temperature have been examined in some coastal areas. At forty-four stations around the British Isles, Dickson (see Wegemann, 1920) found that on an average the diurnal range varied between 0.20 degree in December and 0.69 degree in May. At individual stations both the mean annual range and the variation of the range from month to month were dependent upon the exposure of the locality and the depth of the water at which the measurements were made. This annual variation in range is closely related to the annual variation in the diurnal amount of net heat received by processes of radiation.

Temperature range, °C
Wind and cloudiness Average Maximum Minimum
1. Moderate of fresh breeze
a. Sky Overcast 0.39 0.6 0.0
b. Sky clear 0.71 1.1 0.3
2. Calm or very light breeze
a. Sky overcast 0.93 1.4 0.6
b. Sky clear 1.59 1.9 1.2

The range of the diurnal variation of temperature depends upon the cloudiness and the wind velocity. From observations in the Tropics, Schott (Krümmel, 1907) found the mean and extreme values that are shown in table 32. Similar results but higher numerical values were found by Wegemann from the Challenger data. In both cases the numerical values may be somewhat in error, but the character of the influence of cloudiness and wind is quite evident. With a clear sky the range of the diurnal variation is great, but with great cloudiness it is small, at calm or light breeze it is great, and at moderate or high wind it is small. The effect of cloudiness is explained by the decrease of the diurnal amplitude of the incoming radiation with increasing cloudiness. The effect of the wind is somewhat more complicated, but the main feature is that at high wind velocities the wave motion produces a thorough mixing in the surface layers and the heat which is absorbed in the upper meters is distributed over a thick layer, leading to a small range of the temperature, whereas in calm weather a corresponding intensive mixing does not take place, the heat is not distributed over a thick layer, and consequently the range of the temperature near the surface is much greater.


Diurnal Variation of Temperature in the Upper Layers. Knowledge as to the diurnal variation of temperature at depths below the surface is very scanty. It can be assumed that the depth at which the diurnal variation is perceptible will depend greatly upon the stratification of the water. A sharp increase of density at a short distance below the free-water surface will limit the conduction of heat (p. 477) to such an extent that a diurnal variation of temperature will be present above the boundary surface only.

On the Meteor expedition, hourly temperature observations were made at the surface and at a depth of 50 m at a few stations in the Tropics where an upper homogeneous layer was present which had a thickness of 70 m. Defant (1932) has shown that in these cases the diurnal oscillation of temperature at subsurface depths is in agreement with the laws that have been derived on the assumption of a constant heat conductivity (p. 136). At 50 m the amplitude of the diurnal variation was reduced to less than two tenths of the amplitude at the surface, and the maximum occurred about 6.5 hours later.

The diurnal variation of sea temperature in general is so small that it is of little importance to the physical and biological processes in the sea, but knowledge of the small variations is essential to the study of the diurnal exchange of heat between the atmosphere and the sea (p. 124). The data which are available for this purpose, however, are very inadequate at the present time.

Theory of the Periodic Variations of Temperature at Subsurface Depths

Subsurface temperature variations due to processes of heat conduction can be studied by means of the equation (p. 159)

where ρ is the density and A is the eddy conductivity which, in general, varies with depth and time. When one writes ∂υ/∂t, the local change in temperature, in this form it is supposed that heat conduction takes place in a vertical direction only and that advection can be neglected. The terms “local change” and “advection” are explained on p. 157. An integral of equation (1) is easily found if A/∂ is constant, if the average temperature is a linear function of depth, and if the temperature variations at the surface (z = 0) can be represented by means of a series of harmonic terms:
where σ = 2π/T and T is the period length of the first harmonic term.



Thus, the amplitudes of the harmonic terms decrease exponentially with depth and the phase increases linearly.

Defant (1932) has shown that at the Meteor anchor station no. 288 in latitude 12°38′N, longitude 47°36′W, the diurnal variation of temperature in the upper homogeneous layer indicated a constant eddy conductivity. The amplitudes of the diurnal term at the surface and at 50 m were 0.093 and 0.017 degree, respectively, and the phase difference was 6.5 hours. With ρ = 1.024, and T = 24 hours, he obtained from both the decrease of the amplitude and the difference in phase angle A = 320 g/cm/sec.

In cases in which the annual variation of temperature has been examined, the decrease of amplitude and change of phase give different values of A, indicating that A is not independent of depth and time, as assumed when performing the integration. Fjeldstad (1933) has developed a method for computing the eddy conductivity if it changes with depth, provided that the periodic temperature variations are known at a number of depths between the surface and a depth, h, at which they are supposed to vanish. He arrives at the formula

where an is the amplitude of the nth harmonic term and αn is the phase angle.

Fjeldstad has applied the method to the annual temperature variations off the Bay of Biscay, which have been determined by Helland-Hansen (p. 132). He found, with ρ = 1.025,

Depth (m) 0 25 50 100
Amplitude, a1, °C 3.78 3.24 1.24 0.23
Phase angle, a1 225.1° 235.2° 254.7° 289.3°
Eddy conductivity, g/cm/sec 16.4 3.2 2.1 3.8

Several features show, however, that the observed temperature variations cannot be accounted for by assuming that the eddy conductivity varies with depth only, and variations with seasons also must be considered. Fjeldstad has examined this question and finds that the conductivity reaches a maximum in spring when the stability is at a minimum, but the values remain small throughout the year.

Fjeldstad#x0027;s method can also be applied to the annual variation of temperature in the Kuroshio, which has been discussed by Koenuma (1939) (fig. 32B, p. 131). It is necessary, however, to make the reservation

that in the Kuroshio area the advection term (p. 159) is great (Sverdrup, 1940) and that the use of equation (IV, 5) is therefore correct only if the advection term is independent of time and depth. The harmonic constants and the results, with ρ = 1.025, are

Depth (m) 0 25 50 100 200
Amplitude, a1, °C 4.26 3.97 3.49 2.09 0.71
Amplitude, a2, °C 0.58 0.49 0.44 0.39 0.14
Phase angle, a1 250.2° 253.5° 258.7° 271.8° 289.3°
Phase angle, a2 71.4° 81.0° 100.0° 135.5° 152.6°
Eddy conductivity, A1, g/cm/sec 78 34 23 22 29
Eddy conductivity, A2, g/cm/sec 58 43 39 32 26

Both the annual and the semiannual periods have been used for computing the eddy conductivity, and the agreement between the values of A derived from them must be considered satisfactory, in view of the small amplitudes of the semiannual variation. The numerical values decrease with depth, but are much greater than off the Bay of Biscay, as might be expected, because the high velocity of the Kuroshio must lead to intense turbulence. A possible annual variation of the eddy conductivity has not been examined.

In the Kuroshio region, where the velocity of the current is great and the turbulence correspondingly intense, the annual variation of temperature becomes perceptible to a depth of about 300 m, but in the Bay of Biscay it is very small at 100 m. It can therefore safely be concluded that below a depth of 300 m the temperature of the ocean is not subject to any annual variation.

The eddy conductivity off the Bay of Biscay and in the Kuroshio region is much smaller than that in the upper homogeneous layer near the Equator. The difference can be ascribed to the facts that in the former localities the density increases with depth and that the eddy conductivity is greatly reduced where this takes place (p. 477).

Distribution of Density

The distribution of the density of the ocean waters is characterized by two features. In a vertical direction the stratification is generally stable (p. 416), and in a horizontal direction differences in density can exist only in the presence of currents. The general distribution of density is therefore closely related to the character of the currents, but for the present purposes it is sufficient to emphasize the point that in every ocean region water of a certain density which sinks from the sea surface tends to sink to and spread at depths where that density is found.

Since the density of sea water depends on its temperature and salinity, all processes that alter the temperature or the salinity influence the density. At the surface the density is decreased by heating, addition

of precipitation, melt-water from ice, or run-off from land, and is increased by cooling, evaporation, or formation of ice. If the density of the surface water is increased beyond that of the underlying strata, vertical convection currents arise that lead to the formation of a layer of homogeneous water. Where intensive cooling, evaporation, or freezing takes place, the vertical convection currents penetrate to greater and greater depths until the density has attained a uniform value from the surface to the bottom. When this state has been established, continued increase of the density of the surface water leads to an accumulation of the densest water near the bottom, and, if the process continues in an area which is in free communication with other areas, this bottom water of great density spreads to other regions. Where deep or bottom water of greater density is already present, the sinking water spreads at an intermediate level.

In the open oceans the temperature of the surface water in lower and middle latitudes is so high that the density of the water remains low even in regions where excess evaporation causes high salinities. In these latitudes convection currents are limited to a relatively thin layer near the surface and do not lead to the formation of deep or bottom water. Such formation takes place mainly in high latitudes where, however, the excess of precipitation in most regions prevents the development of convection currents that reach great depths. This excess of precipitation is so great that deep and bottom water is formed only in two cases: (1) if water of high salinity which has been carried into high latitudes by currents is cooled, and (2) if relatively high-salinity water freezes.

The first conditions are encountered in the North Atlantic Ocean where water of the Gulf Stream system, the salinity of which has been raised in lower latitudes by excessive evaporation, is carried into high latitudes. In the Irminger Sea, between Iceland and Greenland, and in the Labrador Sea this water is partly mixed with cold water of low salinity which flows out from the Polar Sea (p. 682). This mixed water has a relatively high salinity, and, when cooled in winter, convection currents that may reach from the surface to the bottom develop before any formation of ice begins. In this manner deep and bottom water is formed which has a high salinity and a temperature which lies several degrees above the freezing point of the water (table 82, p. 683). A similar process takes place in the Norwegian Sea, but there deep and bottom water is formed at a temperature that deviates only slightly from the freezing temperatures (p. 657).

In the Arctic the second process is of minor importance. There the salinity of the surface layers is very low in the regions where freezing occurs, mainly because of the enormous masses of fresh water that are carried into the sea by the Siberian rivers. Close to the Antarctic

Continent formation of bottom water by freezing is of the greatest importance. At some distance from the Antarctic Continent the great excess of precipitation maintains a low surface salinity, and in these areas winter freezing is not great enough to increase the salinity Sufficiently to form bottom water, but on some parts of the continental shelf surrounding Antarctic a rapid freezing in winter leads to the formation of a homogeneous water that attains a higher density than the water off the shelf, and therefore flows down the continental slope to the greatest depths. When sinking, the water is mixed with circumpolar water of somewhat higher temperature and salinity, and hence the resulting bottom water has a temperature slightly above freezing point (p. 611). An active production of bottom water takes place to the south of the Atlantic Ocean, but not within the Antarctic part of the Pacific Ocean.

In some isolated adjacent seas the evaporation may be so intense that a moderate cooling leads to the formation of bottom water. This is the case in the Mediterranean Sea and the Red Sea, and to some extent in the inner part of the Gulf of California, in which the bottom water has a high temperature and salinity and is formed by winter cooling of water whose salinity has been increased greatly by evaporation. Where such seas are in communication with the open oceans, deep water flows out over the sill, mixes with the water masses of the ocean, and spreads out at an intermediate depth corresponding to its density (pp. 670 and 693).

In general, the water of the greatest density is formed in high latitudes, and because this water sinks and fills all ocean basins, the deep and bottom water of all oceans is cold. Only in a few isolated basins in middle latitudes is relatively warm deep and bottom water encountered. When spreading out from the regions of formation the bottom water receives small amounts of heat from the interior of the earth, but this heat is carried away by eddy conduction and currents, and its effect on the temperature distribution is imperceptible.

Sinking of surface water is not limited to regions in which water of particularly high density is formed, but occurs also wherever converging currents (convergences) are present, the sinking water spreading at intermediate depths according to its density. In general, the density of the upper layers increases from the Tropics toward the Poles, and water that sinks at a convergence in a high latitude therefore spreads at a greater depth than water that sinks at a convergence in middle latitudes.

The most conspicuous convergence is the Antarctic Convergence, which can be traced all around the Antarctic Continent (fig. 158, p. 606). The water that sinks at this convergence has a low salinity, but it also has a low temperature and consequently a relatively high density. This water, the Antarctic Intermediate Water, spreads directly over the deep water and is present in all southern oceans at depths between 1200 and

800 m. The corresponding Arctic Convergence is poorly developed in the North Atlantic Ocean, where an Atlantic Arctic Intermediate Water is practically lacking, but is found in the North Pacific, where Pacific Arctic Intermediate Water is typically present.

In middle and lower latitudes two more convergences are found, the Subtropical and the Tropical Convergences. These are not so well defined as the Antarctic Convergence, but must be considered more as regions in which converging currents are present. The Subtropical Convergence is located in latitudes in which the density of the upper layers increases rapidly toward the Poles. The sinking water therefore has a higher density the farther it is removed from the Equator and will spread out at the greater depths.

In the Tropics the density of the surface water is so low that, regardless of how intense a convergence is, water from the surface cannot sink to any appreciable depth but spreads out at a short distance below the surface. A sharp boundary surface develops between this light top layer and the denser water at some greater depth.

In order to account for the general features of the density distribution in the sea, emphasis has been placed on descending motion of surface water, but regions evidently must exist in which ascending motion prevails, because the amount of water that rises toward the surface must exactly equal the amount that sinks. Ascending motion occurs in regions of diverging currents (divergences), which may be present anywhere in the sea but which are particularly conspicuous along the western coasts of the continents, where prevailing winds carry the surface waters away from the coasts. There, the upwelling of subsurface water takes place, which will be described when dealing with specific areas. The upwelling brings water of greater density and lower temperature toward the surface and exercises therefore a widespread influence upon conditions off coasts where the process takes place, but the water rises from depths of less than a few hundred meters. Ascending motion takes place on a large scale around the Antarctic Continent, particularly to the south of the Atlantic Ocean, where rising deep water replaces water that contributes to the formation of the Antarctic bottom water and also replaces water that sinks at the Antarctic Convergence.

It is evident from these considerations that in middle and low latitudes the vertical distribution of density to some extent reflects the horizontal distribution at or near the surface between the Equator and the Poles. It is also evident that, in general, the deeper water in any vertical column is composed of water from different source regions and was once present in the surface somewhere in a higher latitude. Such generalizations are subject, however, to modifications within different ocean areas, owing to the character of the currents, and these modifications will be discussed when dealing with the different oceans.


Subsurface Distribution of Temperature and Salinity

The general distribution of temperature is closely related to that of the density. In high latitudes the temperature is low from the surface to the bottom. The bottom and deep waters that spread out from high latitudes retain their low temperature, but in middle and lower latitudes a warm top layer is present the thickness of which depends partly on the processes of heating and cooling at the surface and partly on the character of the ocean currents. This upper layer of warm water is separated from the deep water by a transition layer within which the temperature decreases rapidly with depth. From analogy with the atmosphere, Defant (1928) has applied the terms troposphere and stratosphere to two different parts of the ocean. Troposphere is applied to the upper layer of relatively high temperature that is found in middle and lower latitudes and within which strong currents are present, and stratosphere to the nearly uniform masses of cold deep and bottom water. This distinction is often a useful one, particularly when dealing with conditions in lower latitudes, but it must be borne in mind that the terms are based on an imperfect analogy with atmospheric conditions and that only some of the characteristics of the atmospheric stratosphere find their counterparts in the sea.

So far we have mainly considered an ideal ocean extending to high northerly and high southerly latitudes. Actually, conditions may be complicated by communication with large basins that contribute to the formation of deep water, such as the Mediterranean Sea, but these cases will be dealt with specifically when we consider the different regions. Conditions will be modified in other directions in the Indian and Pacific Oceans, which are in direct communication with only one of the polar regions, and these modifications will also be taken up later. Here it must be emphasized that the general distribution of temperature is closely related to the distribution of density, which again is controlled by external factors influencing the surface density and the type of deep-sea circulation.

The general distribution of salinity is more complicated than that of temperature. Within the oceanic stratosphere the salinity is very uniform, but within the troposphere it varies greatly, being mainly related to the excess of evaporation over precipitation. The distribution of surface salinity, which was discussed on pp. 124 and 125 is, in general, characteristic of the distribution within the troposphere, as is evident from the vertical section in figs. 210 and 212, which will be dealt with in detail later on.

The Water Masses of the Oceans

TheT-S Diagram. Water masses can be classified on the basis of their temperature-salinity characteristics, but density cannot be used

for classification, because two water masses of different temperatures and salinities may have the same density. For the study of water masses it is convenient to make use of the temperature-salinity diagram, which was introduced by Helland-Hansen (1916). Helland-Hansen points out that when in a given area the temperatures and corresponding salinities of the subsurface water are plotted against each other, the points generally fall on a well-defined curve, the T-S curve, showing the temperature-salinity relationship of the subsurface water of that region. Surface data have to be omitted, because annual variations and local modifications lead to discrepancies.

The corresponding temperature and salinity values in a water column are found to arrange themselves according to depth. The depths of the observed values can be entered along the T-S curve, which then will also give information as to variation of temperature and salinity with depth.


Left: Temperature and salinity at Atlantis stations 1638 and 1640 in the Gulf Stream off Onslow Bay plotted against depth. Right: The same data plotted in a T-S diagram in which σt-curves have been entered.

Since the density of the water at atmospheric pressure, which is expressed by means of σt (p. 56), depends only on temperature and salinity, curves of equal values of σt can be plotted in the T-S diagram. If a sufficiently large scale is used, the exact σt value corresponding to any combination of temperature and salinity can be read off and, if a small scale is used, as is commonly the case, approximate values can be obtained. The slope of the observed T-S curve relative to the σt curves gives immediately an idea of the stability of the stratification (p. 417).

A T-S diagram is shown on the right in fig. 34. On the left in the same figure the observed temperatures and salinities at the Atlantis stations 1638 and 1640 in the Gulf Stream off Onslow Bay are plotted

against depth, and on the right the same values are entered in a T-S diagram. The depths of a few of the observations are indicated. In this case, the temperature-salinity values between 277 and 461 m at station 1638 agree with those between 590 and 790 m at station 1640, indicating that at the two stations water of similar characteristics was present, but at different depths.

The T-S diagram has become one of the most valuable tools in physical oceanography. By means of this diagram characteristic features of the temperature-salinity distribution are conveniently represented and anomalies in the distribution are easily recognized. The diagram is also widely used for detecting possible errors in the determination of temperature or salinity (p. 358).

Water Masses and Their Formation. Following Helland-Hansen's original suggestion, a water mass is defined by a T-S curve, but in exceptional cases it may be defined by a single point in a T-S diagram; that is, by means of a single temperature and a single salinity value. These exceptional cases are encountered in basins where homogeneous water is present over a wide range of depth. A water type, on the other hand, is defined by means of single temperature and salinity values, but a given water type is generally present along a surface in the sea and has no thickness. Only in the exceptional cases that were referred to are the terms “water type” and “water mass” interchangeable, but in oceanographic literature the terms have been used loosely and without the distinction that has been introduced here.

In many areas the T-S curves are straight lines or can be considered as composed of several pieces of straight lines. Elementary considerations show that a linear relation between temperature and salinity must result if the water types that can be defined by the end points of the straight line mix in different proportions. Similarly, a curved T-S relation may result from the mixing of three different types of water. Fig. 35 illustrates in two simple cases how progressive mixing alters the temperature-salinity relation. These considerations are of a formalistic nature, but have in many instances led to the concept that certain water types exist and that the T-S relations that are observed represent the end results of mixing between the types. This concept presupposes that the water types (often referred to as water masses) are continually renewed, because, if that were not the case, processes of mixing would ultimately lead to the formation of homogeneous water. It is possible, however, to account for the character of the T-S curves in the ocean by considering other processes.

In the first place it should be observed that a water mass of uniform temperature and salinity is rarely formed in the open ocean. In high latitudes, where convection currents in winter may reach to the bottom, most of the deep and bottom water will not be uniform, because in some

years the density of the surface water will be greater than in other years and the convection current will reach to different depths, depending upon how much the density of the surface water has been increased. As a consequence, even in these areas the density increases toward the bottom. The bottom water is not homogeneous, and shows therefore a definite temperature-salinity relationship. In the second place, sinking at convergences in middle latitudes may lead, as pointed out by Iselin (1939), to the formation of a water mass with a T-S curve that reflects the horizontal distribution of temperature and salinity at the surface. The upper part of fig. 36 illustrates this point. The figure represents a schematic cross section in which are entered isotherms and isohalines that are all parallel and cut the surface. The σt curves have not been plotted, but are parallel to the isolines. The indicated system will remain stationary if sinking of surface water takes place between the lines a and b and if the sinking water remains on the same σt surface. It will also remain stationary if mixing takes place along or across σt surfaces. These processes will lead to the formation of a water mass that between the curves a and b always shows the same temperature-salinity relation—namely, the relation that is found along the sea surface. Iselin showed that the horizontal T-S curve along the middle part of the North Atlantic Ocean is very similar to the vertical T-S curve that is characteristic between temperatures of 20° and 8° over large areas of the North Atlantic Ocean, and he suggested that processes of sinking and lateral mixing are mainly responsible for the formation of that water. Extensive use of this concept will be made in the chapter dealing with the water masses and currents of the oceans.


Diagrammatic representation of results of vertical mixing of water types. To the left the results of mixing are shown by temperatures and salinities as functions of depth, and to the right are shown in three T-S diagrams the initial water types (1) and the T-S relations produced by progressive mixing (2 and 3).


However, a similar T-S relationship can be established by different processes, as illustrated in the lower part of fig. 36. It is here assumed that two water types, a and b, are formed at the surface and sink along their characteristic σt surfaces. It is furthermore assumed that at subsurface depths mixing takes place between these two water types, whereas near the surface external processes influence the distribution of temperature and salinity so that there the different curves cross each other. In these circumstances one obtains a T-S relation at subsurface depths that is similar to the one found in the previous example, but along the sea surface an entirely different T-S relation may exist. The processes of mixing, in this case, must take place across σt surfaces in order to establish the T-S relation, but at subsurface depths mixing along σt surfaces is not excluded. At present it is impossible to decide what processes are of the greater importance.


Upper: Schematic representation of the formation of a water mass by sinking along σt surfaces (which coincide with the parallel temperature and salinity surfaces) in a region of convergence. The diagram to the right demonstrates that the vertical T-S relation of the water mass agrees with the horizontal T-S relation at the surface in the region of convergence. Lower: Schematic representation of the formation of a water mass by sinking of water types at two convergences and by subsequent mixing. The diagram to the right illustrates that in this case the vertical T-S relation of the water mass need not agree with the horizontal T-S relation at the surface between the convergences.

The point to bear in mind is that the waters of the oceans all attained their original characteristics when the water was in contact with the atmosphere or subject to heating by absorption of radiation in the surface layers, and that in course of time these characteristics may become greatly changed by mixing. This mixing can either be lateral—that

is, take place along σt surfaces, or it can be vertical—that is, crossing σt surfaces.

An example of lateral mixing between water masses is found off the coast of California (Sverdrup and Fleming, 1941), where the water which flows north close to the coast has a T-S relation that differs greatly from that of the water flowing south at some distance from the coast (fig. 199, p. 713). Between these two water masses are found waters of an intermediate character which could not possibly have been formed by vertical mixing and which must have been formed by lateral mixing, probably along σt surfaces. An example of modification of a water mass by vertical mixing is found in the South Atlantic, where the Antarctic intermediate water flows north. This water, near its origin, is characterized by a low salinity minimum, but the greater the distance from the Antarctic Convergence the less pronounced is this minimum (fig. 210, p. 748). The change probably cannot be accounted for by lateral mixing, but Defant (1936) has shown that it can be fully explained as a result of vertical mixing.

Wüst (1935) has introduced a different method for the study of the spreading out and mixing of water types, the “Kernschicht-methode,” which can be translated “the core method.” By the “core” of a layer of water is understood that part of the layer within which temperature or salinity, or both, reach extreme values. Thus, in the Atlantic Ocean, the water that flows out from the Mediterranean has a very high salinity and can be traced over large portions of the Atlantic Ocean by means of a secondary salinity maximum which decreases in intensity with increasing distance from the Strait of Gibraltar. The layer of salinity maximum is considered as the core of the layer in which the Mediterranean water spreads, and the decrease of the salinity within the core is explained as the result of processes of mixing. In this case a certain water type, the Mediterranean water, enters the Atlantic Ocean and loses its characteristic values, owing to the mixing, but can be traced over long distances. The spreading of the water can also be described by means of a T-S curve, one end point of which represents the temperature and salinity at the source region and the other end point of which represents the temperature and salinity in the region where the last trace of this particular water disappears. Having defined such a T-S curve, one can directly read off from the curve the percentage amount of the original water type that is found in any locality. The core method has proved very successful in the Atlantic Ocean and is particularly applicable in cases in which a well-defined water type spreads out from a source region.


Oceanographically a basin is defined as a depression that is filled with sea water and that is partially separated by land or submarine

barriers from the open ocean, with which horizontal communication is restricted to depths less than the greatest depths in the basin. The maximum depth of an entrance from the open ocean to a basin is called the threshold depth, or the sill depth, of that entrance. By “entrance” is understood a depression in a barrier which limits the basin, and it is immaterial whether or not any part of the barrier rises above sea level. The water in the basin is in more or less restricted horizontal communication with the adjacent sea at all levels above that of the lowest sill depth, but below the sill depth renewal of the water in the basin can take place by vertical motion only. It is therefore characteristic of all basins that below the sill depth the water is nearly uniform and approximately of the same character as the water at the sill depth. Other characteristics of the water below sill depth, which will be called the basin water, depend greatly upon the type of exchange of water with the open ocean.

Basins with Outflow Across the Sill. In nearly closed basins in the semiarid regions of lower latitudes, evaporation greatly exceeds precipitation and run-off, and the salinity of the surface water is increased above that of the adjacent open ocean. The evaporation is at a maximum in winter when the surface temperature is simultaneously lowered under the influence of cold continental winds. In winter the surface density is therefore increased so much that vertical convection currents are developed which, in some years when extreme conditions exist, may reach to the greatest depth and bring about renewal of the bottom water. The basin water which is formed in this manner, owing to its very high salinity, will be of greater density than the water at the same depth outside the sill, and must therefore flow out over the sill, following the bottom slope. At some higher level the oceanic water must flow into the basin, partly to compensate for the outflow and partly to compensate for the excess of evaporation over precipitation and run-off. The Mediterranean Sea, the Red Sea, and the inner part of the Gulf of California represent examples of such basins.

In basins of this character the basin water is always characterised by high salinity and generally by high oxygen content. The amounts of inflow and outflow depend upon the difference between evaporation and precipitation and run-off, and the volumes of in- and outflowing water are many times greater than the excess of evaporation. Under stationary conditions the total amount of water which in a given time, Ti, flows into a region must equal the sum of the outflow, Tu, and the difference, D, between evaporation over precipitation and run-off in the same time: Ti = Tu + D. Simultaneously, the amounts of salt carried by the in- and outflowing currents must be equal. In the first approximation (p. 426) formula where formula is the average salinity of the inflowing water and formula is the average salinity of the outflowing water. From these relations one obtains

In basins of this character the inflowing water, which comes from the adjacent open sea, has a relatively high salinity, and therefore the difference, Su − Si, is small. Consequently, the volumes of in- and outflowing water must be great compared to the excess of evaporation over precipitation.

The above considerations are valid only if the entrance is sufficiently wide or deep to permit both inflow and outflow. The Gulf of Kara-Bugaz, on the Caspian Sea, represents an example of a basin which is in such restricted communication with a larger body of water that outflow is practically impossible. This gulf is separated from the Caspian Sea by a 60-mile-long bar and has a shallow entrance which is only a few hundred meters wide. The outflow of the saline deep water is so greatly impeded that, owing to excess evaporation, the salinity of that water in 1902 was 164 ‰ as compared with 12.7 ‰ for the Caspian Sea as a whole.


(A) Basin with local formation of basin water and outflow across the sill. (B) Basin with surface outflow of water of low density and occasional renewal of the basin water by inflow of dense water across the sill.

Conditions are encountered which vary from this extreme case to cases in which the excess evaporation for the entire year is zero, but in which seasonal changes may be large enough to permit occasional development of vertical convection currents reaching to the bottom. The essential features to be emphasized is that in basins of this type renewal of the basin water takes place by vertical convection currents which develop in the basin itself and may reach from surface to bottom. Therefore, the water at and below the sill depth has a higher density than the water at sill depth outside the basin, and is not stagnant.

Basins with Inflow Across the Sill. In the nearly closed basins in higher latitudes, precipitation and run-off exceed evaporation. In such basins a surface layer of low salinity and correspondingly low density is developed. Because of the excess of precipitation and run-off there must be a surface outflow of relatively fresh water, and in order to maintain the salt balance there must be an inflow of more saline water.


The exchange of water with the outside sea is small because the difference, formula, is great. If the difference is so great that the ratio formula is small compared to unity, the relations that are represented by equation (IV, 6) are reduced to

where D now means the excess of precipitation and run-off over evaporation. In these circumstances the inflow is only a small fraction of this excess, and the outflow practically equals the excess.

In basins of this type, stagnant water is often found because renewal of the basin water takes place only if the inflowing water is of greater density than the basin water. Outside the sill the density of the water generally increases much more rapidly with depth than it does inside the sill. Renewal of the basin water takes place if the outside water masses are raised so much that the water which flows in across the sill is of such high density that it sinks toward the bottom of the basin. Fig. 37 shows schematically the character of the exchange with the outside and the renewal of the basin water in the two types of basins.

The rapidity of the renewal of the deep water in the basin depends upon the steepness of the vertical density gradient at the sill depth. If this gradient is steep, an occasional large disturbance fills the basin below sill depth with water of great density, and subsequent disturbances must be as great or greater in order to bring about renewal of the basin water. In extreme cases renewal takes place only by major catastrophes. In the intervals between such catastrophes the basin water may become stagnant, because in the upper layers of stable stratification vertical mixing is insignificant. Some mixing takes place, however, which, between major disturbances, reduces the density of the basin water so much that complete renewal can take place when a new catastrophe occurs.

On the other hand, if the density gradient at the sill depth is small, the outside deep water is brought over the sill by any minor disturbance, and stagnation is prevented by intermittent intrusion of outside deep water, and also by vertical mixing, which is more effective owing to the small density gradient.

The water which sinks at sill depth is heated adiabatically, and the basin water is therefore of nearly constant potential temperature. The effective sill depth—that is, the depth at which the potential temperature in the outside water equals that in the basin—is lower on an average than the actual depth of the sill (table 87, p. 738), and the smaller the density gradient in the outside water, the greater is the difference between the effective and the actual sill depth. Great density gradients, if present, are always found near the surface, and a basin with inflow at

sill depth is therefore likely to have stagnant water if the sill is shallow. The Black Sea, the Baltic, and numerous Norwegian fjords are examples of basins of this type (Fleming and Revelle, 1939, Ström, 1936).

The sill depth has bearing also on the direction of flow across the sill, and this direction does not therefore depend exclusively on an excess or a deficit of evaporation, which was used to facilitate the discussion. At small sill depths an excess or deficit of evaporation determines the character of the exchange, but at great sill depths inflow across the sill develops in most instances. Oceanic water flows freely in and out of the basin at some distance above the sill depth, but at the sill depth the average flow is directed, as a rule, into the basin, because the density of the basin water remains lower than that of the outside water, owing to more effective vertical mixing in a restricted region. The main in- and outflow takes place, however, at lesser depths, the water often flowing in through one entrance and out through another. The basins of the American Mediterranean Sea serve as examples (p. 639).

In large basins in high latitudes, such as the Norwegian Sea and Baffin Bay, deep water is formed locally, owing to freezing or excessive cooling of high-salinity water, although precipitation exceeds evaporation. In such basins, which might be listed as a third type, stagnating water is not found.


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V. Theory of Distribution of Variables in the Sea

Scalar Fields

It cannot be too strongly emphasized that the ocean is three-dimensional and that the distribution of properties or the type of motion must be represented in space. For this purpose a convenient system of coordinates is needed. Any point in the ocean can be designated by means of its geographic latitude and longitude and its depth below sea level, but if one deals with a small area one may consider the surface of the earth within that area as flat, and can introduce ordinary rectangular coordinates with the horizontal axes at sea level and the vertical axis positive downward. By “sea level” is meant not the actual sea level but an ideal sea level, which is defined as a surface along which no component of gravity acts. The difference between the actual and the ideal sea level will be further explained when dealing with the distribution of pressure (p. 406).

The location in the ocean space of any given surface is completely determined if in every latitude and longitude one knows the depth of the surface below the ideal sea level. In a chart this surface can be represented by means of lines of equal depth below sea level (isobaths), which together render it picture of the topography of the surface. Thus, the topography of the sea bottom is shown by isobaths drawn at selected intervals of depth.

The quantities that must be considered when dealing with the sea are either scalars or vectors. A scalar quantity is a physical quantity whose measure is completely described by a number, that depends on the selected system of units. Pressure, temperature, salinity, density, and oxygen content can be mentioned as examples of scalar quantities. A vector is a physical quantity that is completely described by magnitude and direction. The velocity of a particle, the acceleration of a particle, and the forces acting on a particle are examples of vectors.

The magnitude of a vector, such as the numerical value of the velocity of a particle, is a scalar quantity. A vector can be represented by means of its components along the axes of a coordinate system, and these components are scalar quantities.


A continuous fluid is characterized, at every point in the space which it occupies, by a number of different properties. The space distribution of one particular property is called the field of that property. The field is called a scalar field if the property is a scalar quantity, and a vector field if the quantity is a vector. In the ocean there are scalar fields such as the pressure field, the temperature field, and the density field, and there are vector fields such as the velocity field, the acceleration field, and so on.

The term field was first applied to a vector field in order to describe the distribution of electromagnetic forces. Every student of physics has seen the magnetic field of force demonstrated by means of iron filings placed on a card above a magnet, but this experiment brings out only certain characteristics of the field. It shows the direction of the magnetic forces in one single plane, but it does not show the space distribution or the magnitude of the force of the field.

A scalar field is completely represented by means of equiscalar surfaces—that is, surfaces along which the scalar quantity has the same numerical value. The temperature field in the ocean, for instance, would be completely described if one knew exactly the form of the isothermal surfaces, and, similarly, the pressure field would be fully represented if one knew the form of the isobaric surfaces. However, it is impracticable to prepare space models that show the actual configuration of isothermal surfaces or other equiscalar surfaces in the ocean, and it would be impossible to publish such representations. For practical purposes one must select other forms of representation. One widely used method is to show the lines of intersection between equiscalar surfaces and the coordinate surfaces. A chart showing the distribution of temperature at sea level is an example of such representation. In this case the sea level represents one of the principal coordinate surfaces, and the isotherms represent the lines at which the surfaces of equal temperature in the sea intersect the sea surface. Similarly, a chart showing the distribution of temperature at a depth of 1000 m shows the lines along which the isothermal surfaces intersect the 1000-m surface, whereas the temperature distribution in a vertical section shows the lines along which the isothermal surfaces intersect the vertical plane that is under consideration.

A series of horizontal charts of isotherms in surfaces at different distances below sea level give a representation of the temperature field in the ocean, and a series of vertical sections showing isotherms give another representation of the same field.

On the other hand, one can make use of an entirely different method of representation. Instead of showing the lines along which the isothermal surfaces intersect a coordinate surface, one can represent the isothermal surface itself and can show the lines along which the coordinate surfaces at different distances below sea level intersect that surface. Such a

chart would be a chart of the topography of the isothermal surface in question. A series of such topographic charts prepared for a sufficient number of isothermal surfaces, say for every one degree centigrade, would also give a complete representation of the temperature field in the ocean.

These topographic charts would represent charts of absolute topography, because it is assumed that the depths below the ideal sea level are known. This ideal sea level, however, is a fictitious level that cannot be determined by observations, and all measurements have to be made from the actual sea level. In practice, therefore, the topography of a surface in the ocean will not represent the absolute topography but a relative topography referred to the unknown shape of the actual sea surface. In many instances one need not take the difference between absolute and relative topography into account, because it generally amounts to less than 1 m. For instance, it can be neglected when one deals with isothermal surfaces, because the change in temperature on a fraction of a meter is generally negligible. When dealing with the isobaric surfaces, on the other hand, as will be explained in detail when discussing the field of pressure, one must discriminate sharply between absolute and relative topographies.

These matters have been set forth explicitly, because it is essential to bear in mind that one must always consider distribution in space, which can be fully described by means of equiscalar surfaces. These, however, may have highly complicated forms.

The mathematical definition of an equiscalar surface can be written

where s is the scalar quantity under consideration (temperature, pressure, density, and so forth), ∂ s/∂ x · dx is the change of the scalar on the distance dx, ∂ s/∂ y · dy is the change on the distance dy, and ∂ s/∂ z · dz is the change on the distance dz. Along the equiscalar surface the total change must be zero, as expressed by (V, 1).

In a vertical section in the x-z plane the equiscalar curves are similarly defined by

From the latter equation the slope of the equiscalar surfaces in the x-direction is obtained:

Similarly, the slope in the y-direction is

So far, the discussion has dealt with equiscalar surfaces in general. In practice, one may select these surfaces so that there is a constant difference between the value of the variable at any two surfaces. These surfaces are called standard equiscalar surfaces. In the case of temperature, the isothermal surfaces might be selected for every one degree of temperature; in the case of salinity, the isohaline surfaces might be selected for every 0.1 ‰, and so on. These surfaces would divide the space into thin layers characterized by a constant difference of the quantity at the two boundary surfaces of every layer. Such layers are called equiscalar sheets. It should be noted that the scalar is not constant within this sheet but has a constant average value. It is evident that the thickness of these sheets represents the rate at which the scalar varies in a direction at right angles to the equiscalar surfaces. Where the sheets are thin the variation is great, but where the sheets are thick the variation is small. The rate of variation can be represented by means of a vector whose direction is normal to the equiscalar surface and whose magnitude is inversely proportional to the thickness of the sheet. The vector representing the rate of decrease is generally called the gradient (temperature gradient, pressure gradient), and the vector representing the rate of increase is called the ascendant. If the scalar is called s, then the gradient, G, and the ascendant, A, are defined by the equations

Vectors are printed in bold-face type.

If the field is represented by means of a sufficient number of surfaces, these surfaces will completely define the gradients and ascendants that are characteristic of the distribution. Thus, the special vector fields of gradients and ascendants are entirely described by means of systems of equiscalar surfaces, but other vector fields cannot be described in that manner. Vector fields will be dealt with in chapter XII.

Relation between the Distribution of Properties and the Currents in the Sea

Consider any scalar quantity, s (temperature, salinity, pressure, oxygen content, and so on), the distribution of which is continuous in space and time, so that it can be represented as a function of time and the three space coordinates, s = f(t,x,y,z). Let us assume that this scalar quantity can be considered a property of the individual particles of the fluid. A particle in motion after a time dt will be in a new locality,

x + dx, y + dy, z + dz, where the scalar quantity under consideration has the value s + ds = f(t + dt, x + dx, y + dy, z + dz). The property, s, of the individual particle has therefore been changed by the amount ds in the time dt; that is, the time rate of change is ds/dt. This time rate can also be expressed by the characteristics of the field, because, by means of Taylor's expansion, one has
or, since s = f(t,x,y,z),

Dividing by dt and considering that dx/dt, dy/dt and dz/dt represent the components of the velocity, one obtains

The first term on the right-hand side represents the rate of change in a fixed locality—that is, the local change. The last three terms are together called the advection term, because they represent changes that take place in the presence of currents. This relationship is a purely formalistic one and gives no information as to the processes affecting the distribution; it merely states that within a field the individual time change can be considered as composed of two terms: the local time change and the advection.

A few important points can be brought out by means of the above equation: (1) the distribution of any scalar quantity is stationary—that is, independent of time if the local change is zero (∂s/∂ t = 0); (2) the advection terms disappear if there is no motion or if the field is uniform—that is, if either vx = vy = vz = 0 or ∂ s/∂ x = ∂ s/∂ y = ∂ s/∂ z = 0; (3) when the individual change is zero (ds/dt = 0), the local change is equal to the advection but is of opposite sign; (4) if the field of a property is stationary (∂ s/∂ t = 0) and if, further, the individual time change is zero (ds/dt = 0), equation (V, 4) is reduced to

This equation is fulfilled only if the flow is directed along the equiscalar surfaces of the property, as can be seen by comparison with equation (V, l), or by examination of the two-dimensional case.

Distribution of Conservative Concentrations in the Sea

The discussion has so far been of a purely formalistic nature. If one goes a step further and considers the processes that maintain or tend

to alter distributions, it is of advantage to introduce the term concentration to describe any constituent that is present in a measurable amount in a given volume of sea water. Thus, heat content, total salt (which can be represented with sufficient accuracy by salinity), amount of a given compound, and dissolved oxygen or other gases can be given as concentrations, and the same applies to floating organisms. A concentration is a scalar quantity that is continuous in time and space, and the distribution can therefore be represented by means of fields whose characteristics have been dealt with.

The processes that tend to modify the concentrations can be divided into two groups: external processes, which are active only at the boundary surfaces of the fluid, and internal processes, which are active anywhere in the fluid. The external processes are of importance in determining the concentrations at the boundaries, and the internal processes, together with the boundary values, determine the distribution throughout the fluid.

By conservative concentrations are meant concentrations that are altered locally, except at the boundaries, by processes of diffusion and advection only. Heat content and salinity are two outstanding examples of conservative concentrations. Consider a cube the surfaces of which are of unit area and are normal to the coordinate axes. Through the two surfaces that are normal to the x axis, diffusion leads to a transport in unit time of (Ax)1(∂ s/∂ x)1 and (Ax)2(∂ s/∂ x)2, respectively, where both the coefficient, Ax, and the derivative, ∂ s/∂ x, may vary in the x direction. The coefficient of diffusion enters here in the “kinematic” form (p. 470) as A / ρ, where A is the eddy diffusivity, because concentrations have been defined as amounts per unit volume. The difference per unit length of these transports, ∂/∂ x[(Ax/ρ)(∂ s/∂ x)], represents the net change of concentration in the unit volume due to diffusion. In the presence of a current in the x direction, there will also be a net change of concentration due to advection. The concentration that a current of velocity, vx, transports through a unit surface in unit time is equal to svx, and, if this transport changes in the direction of flow, the concentration per unit volume is altered by —∂(svx)/∂ x. Similar considerations are applicable to transport through the other surfaces of the cube, and the combined local change of concentration, therefore, is the sum of terms representing diffusion and advection:

The last term can be written

but in an incompressible fluid the sum of the terms in parentheses is equal to zero (p. 424). Sea water can be considered as incompressible; therefore,
or, in words: Local time change of concentration equals the effect of diffusion minus effect of advection.

Taking equation (V, 4) into consideration, one obtains

or, in words: Individual time change of concentration equals effect of diffusion.

In practice these equations must be greatly simplified. Consider, for example, a two-dimensional system in which the velocity is directed along the x axis, in which diffusion in the x direction can be neglected, and in which it can be assumed that the coefficient of vertical diffusion, A / ρ, is constant. For such a system the condition for a stationary distribution of s, (∂ s/∂ t = 0), is reduced to

This equation has been used by Defant (1929) and Thorade (1931) for studying the character of stationary distributions and by Defant (1936) for computing the ratio A / vx from observed distributions.

As another example, consider a uniform field for which ∂ s/∂ t=ds/dt and assume that Ax = Ay = 0. The above equations are then reduced to

which represents the equation of temperature conduction (p. 135). It may be observed that temperature is not a concentration according to the above definition, but since the temperature is proportional to the heat content, cpϑ, of a unit volume, s can mean temperature.

Other simplifications of the equations can be made, depending upon the nature of the problem under consideration (Montgomery, 1939, Sverdrup, 1939).

Distribution of Nonconservative Concentrations

By nonconservative concentrations are meant primarily concentrations whose distributions are influenced by biological processes besides

those of mixing and transport by currents. For example, the oxygen content is changed by the production of oxygen by plants in the euphotic zone and by the consumption of oxygen by respiratory processes, the phosphate content and that of other plant nutrients are removed from the water when they are utilized by plants and are returned to solution when organic tissues decompose, or the number of organisms of a given species increases or decreases depending upon the relation of the rate of multiplication to the rate at which organisms die off or are consumed.

The local time change of concentration due to biological processes will be called R. Adding this quantity at the right-hand side of equation (V, 5) one can state

In words, local time change of concentration equals effects of diffusion minus advection plus biological processes. This equation can be simplified in the same manner as equations (V, 5) and (V, 6) (Seiwell, 1937, Sverdrup and Fleming, 1941).

The Principle of Dynamic Equilibrium

Experience shows that in a large body of water comparable, say, to the body of water in the Mediterranean Sea, the average conditions do not change from one year to another. The average distribution of temperature remains unaltered year after year, and the same is true as to the average salinity, oxygen content, and contents of minor constituents. If time intervals longer than a year are considered, say ten-year periods, it is probable that even the average number of different species of organisms remains unaltered, provided that the nonaquatic animal, man, does not upset conditions by exterminating certain species and depleting the stock of others. These unchanging conditions represent a state of delicate dynamic equilibrium between factors that always tend to alter the picture in different directions.

In dealing with conservative concentrations, diffusion and advection are at balance except at the sea surface, where external processes contribute toward maintaining the concentration at a certain level. This was illustrated when discussing the general distribution of surface salinity (p. 125), which was shown to depend on two terms, one that represents the external processes of evaporation and precipitation, and one that represents the internal processes of diffusion and advection. Similarly, the surface temperature depends upon heating and cooling by processes of radiation and by exchange with the atmosphere and upon conduction and advection of heat.


In a study of the subsurface distribution of temperature and salinity, it is not necessary to know the processes that maintain the surface values, but it is sufficient to determine these values empirically. If this could be done and if the processes of diffusion and the currents were known, the general distribution of temperature and salinity could be computed. Conversely, if these distributions were known, information as to diffusion and currents could be obtained. In oceanography only the latter method of approach has been employed.

When nonconservative concentrations are dealt with, the principle of a dynamic equilibrium implies that the effects of diffusion, advection, and biological processes cancel. Of the nonconservative concentrations, only the dissolved gases are greatly influenced by the contact with the atmosphere, and other nonconservative concentrations are practically unaltered by external processes.

Application of the principle of dynamic equilibrium can be illustrated by considering the distribution of oxygen. Below the euphotic zone, biological processes that influence the oxygen content always lead to a consumption of oxygen, and the processes of diffusion and advection therefore must lead to a replenishment that exactly balances the consumption. No further conclusions can be drawn. This obvious consideration has been overlooked, however, and some authors have interpreted a layer of minimum oxygen content as a layer of minimum replenishment (Wüst, 1935), while others have considered it a layer of maximum consumption (Wattenberg, 1938).

Conclusions as to the rapidity of consumption (and replenishment) could be drawn from the known distribution of oxygen only if the consumption depended upon the absolute content of oxygen, but the consumption appears to be independent of the oxygen content until this has been reduced to nearly nil (ZoBell, 1940). When all oxygen has been removed, consumption and replenishment must both be zero, and even this obvious conclusion should not be overlooked.

In certain instances a relation may exist between the oxygen distribution and the character of the current. Assume that a nearly horizontal internal boundary exists which separates currents flowing in opposite directions, that diffusion takes place in a vertical direction only, and that the coefficient of diffusion is independent of z. When dynamic equilibrium exists, equation (V, 9) is then reduced to

Since the consumption equals — R and is always positive, the curvature of s is positive when plotted against z. The curvature cannot remain positive at all depths, and therefore it is probable that s, the oxygen content, must be at a minimum near the boundary surface. Thus, a
minimum in the vertical distribution of oxygen may indicate the presence of a boundary surface at which there are no currents, but such a minimum can also develop under entirely different conditions (Seiwell, 1937).

Similar reasoning is valid when dealing with compounds that are removed from the water by organisms for building up their tissues and are returned to solution as metabolic products or by decomposition of organic tissues. A balance is maintained, but in many cases it is not correct to speak of “replenishment” by advection and diffusion, as in the case of oxygen, because the biological processes may lead to a net replenishment, in which case the physical processes must take care of a corresponding removal. Thus, in the deeper layers phosphates and nitrates are added to the water by decomposition of organic matter and removed by diffusion and advection.

When dealing with populations, similar considerations enter. It must be emphasized especially that the number of organisms present in unit volume of water gives no information as to the processes that operate toward changing the number. A small population of diatoms, say, may divide very rapidly without increasing in number, owing to the presence of grazers that consume diatoms. On the other hand, a large population of diatoms may not indicate a rapid production of organic matter, because further growth may be impossible owing to lack of nutrient salts in the water. The terms “population” and “production” have to be clearly defined and kept separate. (“Population” represents concentration, whereas “production” represents one of the processes that alter the concentration.

Another warning appears to be appropriate—namely, a warning against confusion between individual and local changes (p. 157). From the fact that a local population remains unaltered, it cannot be concluded that the population within the water which passes the locality of observation also remains constant—that is, that the individual time change is zero. Similarly, if a sudden change in population is observed in a given locality, it cannot be concluded that the processes which have been active in that locality have led to a rapid growth, because it is equally possible that a new water mass of other characteristics is passing the locality.

If the external influences were clear, if processes of diffusion and advection were known, and if biological and organic chemical processes were fully understood, the distribution of all concentrations could be accounted for. It would then be possible not only to explain the average distribution but also to account for all periodic and apparently random changes. This is the distant goal, but when working toward it one must be fully aware of the limitations of the different methods of approach.

Thus, complete description of the oxygen distribution below the euphotic zone is theoretically possible if the oxygen content in the surface

layers, the processes of diffusion, the currents, and the oxygen-consuming processes of decomposition were known. On the other hand, information as to oxygen consumption can be obtained not only by an examination of the processes of decomposition but also by a computation of the replenishment of oxygen by diffusion and advection. So far, all of our knowledge as to oxygen consumption below the euphotic zone is based on such computations and not on any consideration of biochemical processes.

The dynamic equilibrium, the importance of which has been stressed, exists only insofar as average conditions within a large body of water and over a considerable length of time are concerned. During any part of the day or year the external or internal processes may be subject to periodic or random variation such that at a given moment no equilibrium exists (∂ s/∂ t ≢ 0). At the surface, heating periodically exceeds cooling, and cooling periodically exceeds heating, as a result of which the surface temperature is subjected to diurnal and annual variations that by processes of conduction are transmitted to greater depths. It is possible that longer periods exist which are related to periodic changes in the energy received from the sun, but these long-period variations are of small amplitudes. In many areas, shifts of currents lead to local changes of the temperature which are periodic in character if the shifts are associated with the seasons, or nonperiodic if they are related to apparently random events. In the discussion of the annual variation of temperature (p. 131) the effect of these different processes was illustrated. Similar reasoning is applicable to periodic and random variations of salinity and also to variations of nonconservative properties.

From what has been stated it is evident that in the discussion of the distribution of concentrations in the sea it is as yet impossible to apply a method of deduction based on knowledge of all processes involved in maintaining the distribution. Instead, one has to follow a winding course, discuss processes and their effects whenever possible, discuss actual distributions if such have been determined, and either interpret these distributions by means of knowledge gained from other sources as to acting processes or draw conclusions as to these processes from the distribution. In some instances the processes that maintain the boundary values can be dealt with at considerable length, but otherwise the observed boundary values have to be accepted without attempts at explanation. In all cases, however, it is essential to bear in mind that one is dealing with concentrations in a continuous medium and that general considerations as set forth here are always applicable.


Defant, Albert. 1929. Stabile Lagerung ozeanischer Wasserkörper und dazu gehörige Stromsysteme. Berlin Universität, Institut f. Meereskunde, Veröff., N.F., A. Geogr.-naturwiss, Reihe, Heft, 19, 33 pp., 1929.
Defant, Albert. 1936. “Ausbreitungs- und Vermischungsvorgänge im Antarktischen Bodenstrom und im Subantarktischen Zwischenwasser. Deutsche Atlantische Exped” . Meteor, 1925–1927, Wiss. Erg., Bd. 6, 2 Teil, 2. Lief, p. 55–96, 1936.Montgomery, R. B.1939. “Ein Versuch, den vertikalen und seitlichen Austausch in der Tiefe der Sprungschicht im äquatorialen Atlantischen Ozean zu bestimmen” . Ann. d. Hydrogr. u. Mar. Meteor., p. 242–46, 1939.Seiwell, H. R.1937. “The minimum oxygen concentration in the western basin of the North Atlantic” . Papers in Physical Oceanogr. and Meteorol., v. 5, 24 pp., 1937. Cambridge and Woods Hole, Mass.Sverdrup, H. U.1939. “Lateral mixing in the deep water of the South Atlantic Ocean” . Jour. Marine Research, v. 2, p. 195–207, 1939.Sverdrup, H. U., and R. H. Fleming. 1941. “The waters off the coast of southern California, March to July, 1937” . Scripps Inst. Oceanogr., Univ. Cralifornis, Bull., v. 4, no. 10, p. 261–378, 1941.Thorade, Hermann. 1931. “Strömung und zungenförmige Ausbreitung des Wassers. Gerlands Beitr” . z. Geophys., Bd. 34, Köppen-Bd. 3, p. 57–76, 1931.Wattenberg, Hermann. 1938. “Die Verteilung des Sauerstoffs und des Phosphats im Atlantischen Ozean. Deutsche Atlantische Exped” . Meteor 1925–1927, Wiss. Erg., Bd. 9, 1. Lief, 132 pp., 1938.Wüst, Georg. 1935. “Die Stratosphäre. Deutsche Atlantische Exped” . Meteor 1925–1927, Wiss. Erg., Bd. 6, 1 Teil, 2. Lief, 288 pp., 1935.ZoBell, C. E.1940. “The effect of oxygen tension on the rate of oxidation of organic matter in sea water by bacteria” . Jour. Marine Research, v. 3, no. 3, p. 211–223, 1940.

VI. Chemistry of Sea Water

If suspended solid material of either organic or inorganic origin is excluded, sea water may be considered as an aqueous solution containing a variety of dissolved solids and gases. Determination of the chemical nature and concentrations of the dissolved substances is difficult for the following reasons: (1) some of the dissolved substances, such as chloride and sodium ions, are present in very high concentrations, while others, certain metals for instance, are present in such minute quantities that they have not been detected in sea water, although they have been found in marine organisms or salt deposits; (2) two of the major constituents, sodium and potassium, are extremely difficult to determine accurately; (3) it is virtually impossible in some cases to separate related substances such as phosphate and arsenate, calcium and strontium, and chloride, bromide, and iodide. In these cases the combined elements are determined together and usually reported as if they represented only one; that is, calcium and strontium are often calculated as “calcium,” and chloride, bromide, and iodide as “chloride.”

Because of the complex nature of the dissolved materials in sea water a specially developed technique is usually required to determine the concentration of any constituent. The standard methods for the quantitative analysis of solutions which are given in textbooks generally cannot be applied to sea water without adequate checks on their accuracy. This is particularly true when dealing with elements present in extremely low concentrations, because the elements occurring as impurities in the reagents may be in amounts many times those found in the water.

Constancy of Composition

It has been found that, regardless of the absolute concentration of the total solids, the ratios between the more abundant substances are virtually constant. The importance of this result cannot be overemphasized, as upon it depends the validity of the chlorinity:salinity:density relationships and, hence, the accuracy of all conclusions based on the distribution of density where the latter is determined by chemical or indirect physical methods such as electrical conductivity or refractive index.


The relative uniformity in the composition of the sea water was established by the investigations of Forchhammer, Natterer, and Dittmar. Although Forchhammer analyzed a large number of samples, his investigations were not complete because he did not determine certain of the abundant elements. Natterer made more detailed analyses, but it was Dittmar who laid the solid foundation for the present knowledge of the composition of sea water.

Dittmar (1884) made careful determinations on 77 water samples, representative of all oceans, which had been collected on the voyage around the world of H.M.S. Challenger. He determined the halides, sulphate, magnesium, calcium, and potassium. On composite samples he found the ratio of bromine to chlorine and estimated the carbonate. From the sums of the chemical equivalents of the negative and positive ions, he calculated the sodium by difference. This procedure was followed because he was unable to achieve satisfactory direct determinations for sodium. The results of Dittmar's work showed that there were no significant regional differences in the relative composition of sea water; consequently his average values could be used to represent the ratios between the major dissolved constituents. In table 33 are given Dittmar's average values in the units in use at the present time and referred to a chlorinity of 19.00 ‰. The percentages of the various ions are also shown.

Ion Original values Recalculated, 1940 atomic weights 1940 values
Cl = 19 ‰ % Cl = 19 ‰ % Cl = 19 ‰ %
C1 18.971 55.29 18.971 55.26 18.980 55.04
Br 0.065 0.19 0.065 0.19 0.065 0.19
SO42.639 7.69 2.635 7.68 2.649 7.68
CO3 0.071 0.21 0.071 0.21 …… ……
HCO3 …… …… …… …… 0.140 0.41
F …… …… …… …… 0.001 0.00
H3BO3 …… …… …… …… 0.026 0.07
Mg++ 1.278 3.72 1.292 3.76 1.272 3.69
Ca++ 0.411 1.20 0.411 1.20 0.400 1.16
Sr++ 0.013 0.04
K+ 0.379 1.10 0.385 1.12 0.380 1.10
Na+ 10.497 30.59 10.498 30.58 10.556 30.61
Total 34.311 34.328 34.482


Since 1884 the modification of atomic weights has affected the numerical results reported by Dittmar. Corrections for these changes may be made (Lyman and Fleming, 1940) as shown in the “recalculated” values in table 33. In the latter tabulation the sodium has been recalculated by Difference.

It is interesting to compare Dittmar's results with those obtained by modern methods of analysis as shown in the last columns of the table. The sources of these data are indicated in table 35. It is immediately seen that there are small differences for most of the elements determined by Dittmar and that certain other ions have been added to the list of major constituents. The bound carbon dioxide is reported as bicarbonate ion instead of as carbonate, strontium is given by itself instead of in combination with calcium, and fluoride and boric acid have been added.

The close agreement between the results of Dittmar and those obtained recently is remarkable when we consider the complexity of the problem and the great advance in knowledge of analytical chemistry. However, although the differences are small, they are significant, and hence the importance of Dittmar's work is that it showed the constancy of the ratios between the major constituents, and not that it led to accurate numerical values of these ratios.

In table 33 the composition is shown by referring the substances to a standard concentration, C1 = 19.00 ‰, and by means of the ratios between the different ions and the total dissolved solids. In most instances it is preferable to use a third method; namely, to give the ratios between the various substances and the chlorinity or the chlorosity (p. 52), and these ratios are known as C1-ratios and chlorosity factors, respectively. The Cl-ratio is the amount of any ion or substance per unit (gram) of chlorinity, and is obtained by dividing the concentration in grams per kilogram by the chlorinity, or the concentration in grams per 20°-liter by the chlorosity. Multiplication of the C1-ratio by a given chlorinity or corresponding chlorosity will give the concentrations as grams per kilogram or per liter, respectively. Concentrations in milligram-atom units are always on a liter basis, and, if divided by the chlorosity, yield the ratios that are called chlorosity factors. It may be noted that a chlorosity factor multiplied by chlorinity yields the concentration in milligram-atoms per kilogram.

The uniformity of relative composition in the oceans is the result of circulation and mixing. These operations are continuous, and tend to eliminate regional differences in composition, whatever the cause. Disturbing agencies bring about changes that are small compared to the bulk of the substances present and consequently will not materially affect the relative concentration of the major constituents. Further-more, many of the disturbing processes that tend to modify the relative composition are reversible. For example, the secretion of calcium

carbonate by organisms, which reduces the quantity of calcium in solution, takes place at a certain season or in certain parts of the sea, but upon the death of the organisms the calcium carbonate may dissolve in other regions. Although small regional differences may result from such processes, the effects are largely neutralized by mixing. On the basis of parallel determinations of density by means of interferometer (p. 54) and chlorine titration, Lotte Möller (Bein, Hirsekorn, and Möller, 1935) has shown that very small systematic differences exist in the composition of water masses of the North Atlantic, but as yet these are significant only as refined means for tracing water masses of certain characteristics.

The constancy of composition is, as already emphasized, of the greatest importance. Not only is it the basis of the chlorinity:salinity:density relationships, but it also affords a means of estimating the concentrations of all of the major constituents when the concentration of any one of them is known. Furthermore, results of studies on the composition or the physical properties of sea water in any locality are generally applicable to the water in any other part of the oceans.

Except in special areas, such as in the Baltic Sea, the Black Sea, and off the mouths of large rivers, it is not necessary to consider that the water represents special local types with properties that differ from those of sea water in general. Nevertheless, it should be remembered that the composition is not absolutely constant even for the major constituents listed in table 33. Various factors which will be discussed in detail later are always operating and always tend to modify the relative abundances. Rivers introduce dissolved material in proportions that are markedly different from those in the sea, and they also introduce sedimentary material that reacts in various ways with the dissolved constituents. The formation and melting of sea ice may bring about a modified distribution of the dissolved substances.

Thus far, comment has been largely restricted to those constituents of sea water that are present in large, or at least relatively constant, proportions. If we consider those elements which are present in small quantities and which are utilized by marine organisms, the concept of constant composition is no longer generally valid, because the concentrations of these elements vary widely, particularly near the surface. A great part of the work in chemical oceanography is now devoted to determining the space and time variations in variable constituents, and much thought is directed toward the solution of the problems related to the processes that control the observed distribution.

Units Used in Chemical Oceanography

In chemical oceanography most of the numerical results are expressed as concentrations—that is, as the amounts of various constituents in a certain quantity of sea water. Obviously many different combinations

of mass and volume units are possible and, in fact, a great variety have been used. In order to avoid confusion and to make the results of different workers directly comparable, it is desirable that a standardized system of units for reporting results in chemical oceanography be followed. Such a system has been proposed by the International Association of Physical Oceanography (1939).

Only two units are to be used for expressing the quantity of sea water: either (1) the kilogram or (2) the amount of water which at 20° C. and pressure one atmosphere occupies the volume of one liter. The latter unit is designated as L20, but in this discussion it will be indicated as L. The system in which the constituents are reported as the amounts present per liter is designated as the “preferred” one, with an alternative for the abundant substances that may be reported as grams per kilogram of sea water. Salinity and chlorinity are always reported as grams per kilogram of sea water. It should be understood that the proposed system applies only to the reporting of analytical data in the literature. Any suitable units may be adopted for the discussion of special problems.

For expressing the amounts of the dissolved constituents, two types of units are proposed: (1) physical units of mass, volume, or pressure, and (2) units based upon the number of atoms of the designated element, which may be present as ions or molecules either singly or in combination with other elements. In certain cases the number of chemical equivalents is acceptable.

The mass units most commonly used are those of the metric system and bear the following relations to each other:

A measure of the number of atoms of the designated element is obtained by dividing the amount of the element, expressed as grams, milligrams, or mygrams, by the gram-atomic weight of the element. Hence,
Quantities expressed as gram-, milligram-, or mygram-atoms may be converted to the corresponding mass units by multiplying by the gram-atomic weight of the designated element.

In certain cases (for example, alkalinity and hydrogen-ion concentration) it is desirable to report the concentration in terms of chemical equivalents. The units shall then be



For expressing the partial pressure of gases dissolved in sea water the basic pressure unit is the “physical atmosphere” (p. 55):

Partial pressures shall be expressed in Torr.

Volume units are all based upon the true liter—that is, the volume of 1 kg of distilled water at 4°C. When volume units are used, the temperature and pressure should be stated. The quantities of dissolved gases, when expressed as milliliters (ml), should be those for 0°C and a pressure of 1 atmosphere, that is, NTP.

The centigrade scale is to be used for reporting temperatures.

The units to be used in reporting data, proposed by the International Association of Physical Oceanography, are given in table 34. It should be noted that all units are based upon the amount of a designated element that may be present either singly (for example, oxygen or calcium) or in combination with other elements (for example, phosphate-phosphorus).

Because the 20° liter is the standard volume unit for expressing the quantity of sea water, glassware should be calibrated for this temperature, and, if practicable, measurements and chemical determinations should be made at or near this temperature. If the sea-water samples are not at 20°, it may be necessary to apply certain corrections. Full descriptions of the methods for making such corrections and tables to facilitate the transformation are included in the Report of the International Association of Physical Oceanography. In most cases the accuracy of the methods of analysis for the elements present in small amounts do not justify such corrections.

As already stated, it is frequently desirable to express the relative concentrations as Cl-ratios or chlorosity factors (p. 167). These relationships may be used to calculate the quantity of the major elements present in water of known chlorinity or to check variations in composition which may be brought about by natural agencies, pollution by sewage and industrial wastes, or by other agencies.

Composition of Sea Water

So far, the discussion of the composition of sea water has been based mainly on the results of the fundamental investigations of Dittmar. Since his time our knowledge of the composition of sea water has increased tremendously. Improved methods of analysis have been developed and consequently more accurate values can be obtained. Tests have also been developed for the detection and determination of elements other than those previously discussed. Particular efforts have been devoted to the study of the so-called plant nutrients—that is, those elements

which are essential to plant growth in the sea but which are present in small and variable amounts. Individual elements have been studied both extensively and intensively, so that much more is now known concerning the regional differences in the ratios of the major constituents and in the amounts of the elements present in small quantities. However, except for Dittmar's work, there has been no careful study of the composition of a large number of samples for all the major constituents.

ABBREVIATIONS AND UNITS TO BE USED IN REPORTING CHEMICAL DATA (Scheme proposed by the International Association of Physical Oceanography)
Designated substance Abbreviation Units (p = preferred, a = alternative)




Ammonia-nitrogen Ammonia-N p
Argon Argon p
Arsenate-arsenic Arsenate-As p
Arsenite-arsenic Arsenite-As p
Borate-boron Borate-B p
Calcium Ca p a
Carbon dioxide Carbon dioxide-C p
CO2 a
Chlorinity Cl p
Copper Cu p
Iron Fe p
Magnesium Mg p a
Manganese Mn p
Nitrate-nitrogen Nitrate-N p
Nitrite-nitrogen Nitrite-N p
Nitrogen (gas) N2 p a
Oxygen (gas) O2 p a
Phosphate-phosphorus Phosphate-P p
Potassium K p a
Radioactive substances p
Salinity S p
Silicate-silicon Silicate-Si p
Sodium Na p a
Sulphate Sulphate-S p
SO4 a
Hydrogen sulphide Sulphide-S p
H2S a


Hence, in order to prepare a tabulation of the composition of sea water it is necessary to combine the results of numerous workers who have examined samples from different sources. All available data were collected by Thompson and Robinson (1932), and additional references will be found in the following discussion. In some cases the information is extensive, but for other elements only a few determinations have been made on water from a single locality. We shall first examine the quantities of the major elements—that is, those which bear a virtually constant relationship to the chlorinity.

In table 35 is given a compilation of the major ions that make up over 99.9 per cent of the known dissolved solid constituents of sea water. The sources of these data have been discussed by Lyman and Fleming (1940). The concentrations of the various ions are shown for water of 19.00 ‰ chlorinity, and also the Cl-ratios. The quantities are also expressed in terms of chemical equivalents per kilogram for water of 19.00 ‰ chlorinity and as milligram-atoms per 20° liter. Chlorosity factors are given for units of milligram-atoms. The carbon dioxide has been reported as bicarbonate. This method is not strictly accurate, because the bound carbon dioxide content of sea water is variable, but, as will be shown in the discussion of the carbon dioxide system, the sum of the chemical equivalents of carbonate and bicarbonate is virtually constant for any chlorinity.

It is immediately seen that the sum of the halides (chloride, bromide, and fluoride) by weight is greater than the chlorinity. The amount of iodide is negligible. Even if the bromide is calculated as chloride, and if the fluoride is disregarded because it does not take part in the chlorinity determination, the chloride equivalent is 1.00045 times greater than the chlorinity. The reasons for this apparent discrepancy have been discussed on page 52.

Lyman and Fleming (1940) obtained the following empirical equation for the dissolved solids as represented in table 35:


From this it will be seen that in water of 19.00 ‰ chlorinity the total dissolved solids are 34.4816 ‰, but, according to the equation used to calculate the salinity from the chlorinity (p. 51), the salinity is 34.325 ‰. Thus, the total amount of dissolved solids is greater than the salinity. If, on the other hand, the salinity is calculated from the total solids, using the definition for the former quantity—that is, by converting the bicarbonate to oxide and converting the bromide to chloride—we obtain the salinity “by definition” as 34.324 ‰. This agreement must be considered as more or less accidental, as there are many uncertainties in the analytical data. Confidence in the values is strengthened, however, by the fact that the sodium: chlorinity ratio as reported by Robinson and Knapman (1941) agrees exactly with the value that Lyman and Fleming (1940) found by difference. Although the table represents the most probable figures for the composition of the major dissolved constituents, it is subject to change as additional data become available.

MAJOR CONSTITUENTS OF SEA WATER(Cl = 19.00 ‰, ρ20 = 1.0243)
Ion Cl-ratio, g per unit Cl Equivalent per kg of sea water mg-atoms per liter Chlorosity factor, mg-atoms per unit Cl Authority

Total dissolved solids = 34.4816 ‰

Sum of constituents (HCO3− as O, and Br as Cl) = 34.324 ‰

Salinity (S ‰ = 0.030 + 1.805 Cl ‰) = 34.325 ‰

aRatio for millival/kg = 0.1205

bRatio for boron/Cl = 0.000240

cBoric acid undissociated

dSodium calculated by difference in sum of equivalents

Chloride, Cl 18.9799 0.99894 0.5353 548.30 28.173 Dittmar (1884), Jacobsen and Knudsen (1940)
Sulphate, SO4 2.6486 0.1394 0.0551 (SO4-S) 38.24 1.451 Thompson, Johnston, and Wirth (1931)
Bicarbonate, HC03 0.1397 0.00735[a] 0.0023 (HCO3C) 2.34 0.120 Revelle (1936)
Bromide, Br 0.0646 0.00340 0.0008 0.83 0.0426 Dittmar (1884)
Fluoride, F 0.0013 0.00007 0.0001 0.07 0.0036 Thompson and Taylor (1933)
Boric acid,c H3BO3 0.0260 0.00137[b] [c] (H3BO3-B)0.43 0.0221 Harding and Moberg (1934), Igelsrud, Thompson, and Zwicker (1938)
Total 0.5936
Sodium,[d] Na+ 10.5561 0.5556 0.4590 470.15 24.153 By difference, and Robinson and Knapman (1941)
Magnesium, Mg++ 1.2720 0.06695 0.1046 53.57 2.752 Thompson and Wright (1930)
Calcium, Ca++ 0.4001 0.02106 0.0200 10.24 0.5262 Kirk and Moberg (1933); Thompson and Wright (1930)
Potassium, K+ 0.3800 0.02000 0.0097 9.96 0.5113 Thompson and Robinson (1932)
Strontium, Sr++ 0.0133 0.00070 0.0003 0.15 0.0077 Webb (1938)
Total 0.5936


The data in table 35 apply more specifically to surface water than to deep water. Both bicarbonate ion and calcium will be slightly higher in deeper water. Furthermore, some of the other compounds not included in this tabulation, such as nitrate and silicate, may be present in sufficient quantities to disturb the balance of the anions and cations shown in the table. The Cl-ratios should therefore be considered more as indices than as absolute values. However, in no case will the ratios vary by more than a unit or two in the last decimal place when the water under investigation is from the open sea. Under abnormal conditions, as in highly diluted water, larger departures may be found. By definition the salinity is not zero at zero chlorinity; hence the ratios of certain elements would be expected to approach infinity at very high dilutions when the diluting water contained substances other than halides. Therefore, in studies in areas of highly diluted water the character of the river water should be taken into account. As pollution problems frequently occur in such areas, it will be necessary to determine the normal ratios for different dilutions for a specific zone before any conclusions can be drawn as to the nature or extent of the pollution.

Elements Present in Sea Water

Thus far, only the major constituents of sea water have been considered. In table 36 are entered all elements that are known to occur in sea water as dissolved solids, except hydrogen and oxygen. They are not given as ions in this case but as the amounts of the individual elements which occur in water of chlorinity 19.00 ‰. The elements are arranged in the order of their abundance. In the first column they are reported as milligrams per kilogram, and in the second as milligram-atoms per liter. For convenience, the 1940 atomic weights and their reciprocals have been included. These constants are necessary when converting weight units to gram-atom units, and vice versa. The values for the major elements correspond to those given in previous tables and, in general, are valid for surface water. For many of the elements ranges in concentration have been indicated. No doubt ranges should be shown for others, but the lack of sufficient observations or uncertainty as to the reliability of reported data leaves these problems unsettled. For many of the elements that are present in very low concentrations there are only one or two determinations available, and in some cases only indirect estimates have been made. Hence, in these cases the indicated values can represent only the order of magnitude of the quantities present. Omitting the six most abundant elements, only carbon (CO2 components), silicon,

nitrogen, and phosphorus compounds have been studied with sufficient completeness to provide a fairly good idea of their distribution. Less complete studies have been made on the variations in the amounts of boron, iodine, iron, manganese, copper, gold, and radium. Cadmium, chromium, cobalt, and tin have been found in the ash of marine organisms, and hence it is implied that they occur in sea water, although so far they have not been shown directly.

Forty-four elements are listed in table 36, and if we add hydrogen, oxygen, and the inert gases neon, helium, and argon, we obtain a total of forty-nine elements that are known to occur in sea water. Further investigations will undoubtedly demonstrate the presence of others. Certain problems of the origin and concentration of the dissolved solids relative to their concentration in the earth's crust will be discussed later.

The following brief discussion is limited to those elements that either occur in relatively large amounts or whose distribution has been shown to be affected by biological activity. For elements in the latter group additional data are given in chapter VII. In table 36 references are given for those elements not discussed in the text. A comprehensive discussion is given by Thompson and Robinson (1932), and other results are reported by Goldschmidt (1937) and Wattenberg (1938). The elements are considered in the order in which they appear in the table.

Chlorine, present as chloride ion, is the most abundant ion and makes up about 55 per cent by weight of the dissolved material. It is rarely measured except in combination with other halides in the chlorinity determination. The bromide and iodide are then computed as if they were chloride. It should be kept in mind that the ratio of the chlorine equivalent of the halides to the chlorinity is 1.00045 (p. 52). The chlorinity is of the greatest importance, not only as the basis of density computations, but also as the standard to which those substances present in major amounts are referred.

Sodium is the most abundant cation in sea water, but it is rarely determined directly, owing to the technical difficulties involved in the determination of the alkali metals. The average ratio to chlorinity, 0.5556, as obtained by Robinson and Knapman (1941) agrees exactly with the value that Lyman and Fleming (1940) calculated by difference. It is somewhat higher than the average of 0.5509 given by Thompson and Robinson (1932), but is in fair agreement with the ratio 0.5549 obtained by Webb (1939) by direct analysis. The sodium: chlorinity ratio may be modified near river mouths.

Magnesium content of sea water has been investigated rather carefully, particularly by Thompson and Wright (1930). The magnesium is usually determined by a special modification of the magnesium-ammonium-phosphate method. The ratio of magnesium to chlorinity is very uniform.

ELEMENTS PRESENT IN SOLUTION IN SEA WATER (Dissolved gases not included)
Element mg/kg Cl = 19.00 ‰ mg-atoms/L Cl = 19.00 ‰ Atomic weight (1940) 1/atomic weight Authority
Chlorine 18980 548.30 35.457 0.02820
Sodium 10561 470.15 22.997 0.04348
Magnesium 1272 53.57 24.32 0.04112
Sulphur 884 28.24 32.06 0.03119
Calcium 400 10.24 40.08 0.02495
Potassium 380 9.96 39.096 0.02558
Bromine 65 0.83 79.916 0.01251
Carbon 28 2.34 12.01 0.08326
Strontium 13 0.15 87.63 0.01141
Boron 4.6 0.43 10.82 0.09242
Silicon 0.02 –4.0 0.0007 –0.14 28.06 0.03564
Fluorine 1.4 0.07 19.00 0.05263
Nitrogen (comp.) 0.01 –0.7 0.001 –0.05 14.008 0.07139
Aluminum 0.5 0.02 26.97 0.03708
Rubidium 0.2 0.002 85.48 0.01170
Lithium 0.1 0.014 6.940 0.14409
Phosphorus 0.001–0.10 0.00003–0.003 30.98 0.03228
Barium 0.05 0.0004 137.36 0.00728
Iodine 0.05 0.0004 126.92 0.00788
Arsenic 0.01–0.02 0.00015–0.0003 74.91 0.01335
Iron 0.002–0.02 0.00003–0.0003 55.85 0.01791
Manganese 0.001–0.01 0.00002–0.0002 54.93 0.01820
Copper 0.001–0.01 0.00002–0.0002 63.57 0.01573
Zinc 0.005 0.00008 65.38 0.01530 Atkins (1936)
Lead 0.004 0.00002 207.21 0.00483 Boury (1938)
Selenium 0.004 0.00005 78.96 0.01266 Goldschmidt and Strock (1935)


0.002 0.00002 132.91 0.00752 Wattenberg (1938)
Uranium 0.0015 0.00001 238.07 0.00420 Föyn et al (1939)
Molybdenum 0.0005 0.000005 95.95 0.01042 Ernst and Hoermann (1936)
Thorium <0.0005 <0.000002 232.12 0.00431 Föyn et al (1939)
Cerium 0.0004 0.000003 140.13 0.00714 Goldschmidt (1937)
Silver 0.0003 0.000003 107.880 0.00927 Haber (1928)
Vanadium 0.0003 0.000006 50.95 0.01963 Ernst and Hoermann (1936)
Lanthanum 0.0003 0.000002 138.92 0.00720 Goldschmidt (1937)
Yttrium 0.0003 0.000003 88.92 0.01125 Goldschmidt (1937)
Nickel 0.0001 0.000002 58.69 0.01704 Ernst and Hoermann (1936)
Scandium 0.00004 0.0000009 45.10 0.02217 Goldschmidt (1937)
Mercury 0.00003 0.0000001 200.61 0.00498 Goldschmidt (1937)
Gold 0.000006 0.00000002 197.2 0.00507 Haber (1928)
Radium 0.2 – 3 × 10−10 0.8 – 12 × 10−13 226.05 0.00442 Evans, Kip, and Moberg (1938)
Cadmium Fox and Ramage (1931)
Chromium Webb (1937)
Cobalt Thompson and Robinson (1932)
Tin Thompson and Robinson (1932)


Sulphur is present in sea water as sulphate ion, and is in this form usually determined by precipitation as barium sulphate. An extensive study of the sulphate distribution has been made by Thompson, Johnston, and Wirth (1931). Under stagnant conditions occurring in certain isolated basins, and in and near bottom sediments, a part of the sulphate may be converted to sulphide ion. Considerable quantities of sulphide occur in the Black Sea and in certain Norwegian fjords, and its presence has been reported in many localities. The sulphate: chlorinity ratio may also be modified by dilution with river water, which is generally relatively high in sulphate. Processes of freezing and melting may possibly affect the relative concentration (p. 216).

Calcium is present in much smaller quantities than either sodium or magnesium, but its distribution in the ocean has been studied much more thoroughly, mainly because calcium is a major constituent of many skeletal remains found in marine sediments. By deposition of such remains calcium is permanently removed from the water, but this removal does not necessarily imply that the calcium concentration is decreasing, because a large supply is maintained by the river waters flowing into the sea. Detectable differences in the calcium: chlorinity ratio have been observed. In the Baltic, Gripenberg (1937a) has shown that the type of river water which has diluted the sea water can be determined from that ratio. Furthermore, Moberg and Revelle (1937) have demonstrated the existence of vertical differences in the calcium: chlorinity ratio which they attribute to the removal of calcium in the surface layers through biological activity. Interest in the concentration of calcium has also centered around the question of the solubility of calcium carbonate in sea water and the factors that control precipitation and solution. In certain areas calcium carbonate is apparently precipitated inorganically, and in other regions it apparently passes into solution. In addition to these problems, knowledge of the calcium concentration is important in an understanding of the carbon dioxide system in the sea, which will be discussed later. The quantity of calcium is usually determined by precipitation as the oxalate under carefully controlled conditions and subsequent titration with potassium permanganate. One such method has been described by Kirk and Moberg (1933).

Webb has pointed out that in this method for the estimation of calcium the strontium will be carried down, and hence the calcium figure will be too high by the equivalent amount of strontium. As the ratio calcium: strontium is apparently constant, Webb suggests that the “calcium” shall be taken to mean the calcium after the strontium and barium have been replaced by calcium. Since the barium is negligible in this case, the values of “calcium” will be given directly by volumetric methods, but when the quantities are determined by weighing, corrections

must be applied (Webb, 1938). Values cited in this discussion are for calcium alone and have been obtained by correcting the analytical data for the presence of strontium. The “calcium” Cl-ratio as defined by Webb and corresponding to the values of calcium and strontium in table 35 is 0.0214.

Potassium is the fourth most abundant cation and is present in amounts of only a few per cent of that of sodium. The potassium is rarely determined directly, but apparently it bears a very constant relationship to the chlorinity (Thompson and Robinson, 1932). However, the content of potassium may be modified by biological agencies, since some organisms, particularly the large algae, concentrate potassium to a marked degree. The ratio of the potassium to chlorinity may also be modified by dilution with river water. The potassium may react with the colloidal and clay particles brought to the sea by rivers and run-off, and consequently this agency may influence the ratio. Certain minerals formed on the sea bottom, such as glauconite, contain potassium.

Bromine shows a very constant ratio to the chlorinity and is apparently all present as bromide ion.

Discussion of the concentration of carbon in sea water is complicated by the fact that it occurs not only in the form of carbonic acid and its salts but also in appreciable amounts as a constituent of organic material, either living or dead. The detrital organic material may be either particulate or in solution. The solubility of carbon dioxide depends upon the temperature and salinity of the water, and exchange of carbon dioxide with the atmosphere takes place at the surface. Photosynthesis in the surface layers reduces the amount of carbon dioxide in the water, and respiration increases the concentration. Consequently, the quantities of carbon present as either free carbon dioxide, bicarbonate, or carbonate will show a considerable range. These problems will be discussed in the sections dealing with the carbon dioxide system in the sea. The quantity of carbon given in table 36 was calculated on the assumption that only bicarbonate ions were present. The organic carbon, which is probably of the order of 2 to 3 mg/L (0.15 to 0.25 mg-atoms/L), was not included. The methods by which the different carbon dioxide components and the particulate and dissolved organic carbon may be determined are discussed later.

Strontium has not been investigated in detail, as it is extremely difficult to determine quantitatively. In determinations of calcium by means of the oxalate precipitation, the strontium is carried down with the calcium, and consequently the ratio of calcium: chlorinity usually reported for sea water represents the calcium plus strontium reported as calcium. Strontium is a constituent of the calcareous skeletons of certain organisms.


Boron occurs in sea water in a surprisingly high concentration and bears a constant relationship to the chlorinity. Apparently it is present as undissociated boric acid. There has been considerable uncertainty as to the form in which boron occurs, but the method of determination is standardized against boric acid and the values can at least be expressed as equivalent to a certain concentration of boric acid. The determination of boric acid in sea water is based on titration with very dilute sodium hydroxide in the presence of mannitol. Methods have been described by Harding and Moberg (1934) and by Igelsrud, Thompson, and Zwicker (1938). The amount of boron present in sea water is of interest in the carbonate equilibria and in this connection will be discussed later. Boron is concentrated by certain marine organisms.

Silicon has been studied extensively because it is utilized by diatoms and other silica-secreting organisms. According to a tabulation by Thompson and Robinson (1932), the silicate-silicon varies by more than one hundredfold—namely, from 0.0007 to 0.11 mg-atoms/L (0.02 to 3.0 mg/L). Clowes (1938) found values slightly exceeding .14 mg-atoms/L (4.0 mg/L) in the deep waters of the Antarctic. Surface samples are usually low, owing to the development of silica-secreting organisms, but a progressive increase in silicate takes place with depth, which is ascribed to the dissolving of soluble silicates. However, there is always the possibility that the water contains silicon in some compound present in colloidal form. River water contains a high content of silicon, both in solution and as colloidal particles. Diatom and radiolarian oozes contain the siliceous remains of organisms that have developed near the surface and settled to the bottom after their death. Although siliceous deposits of organic origin cover large areas, most of the siliceous skeletal remains dissolve after the death of the organisms. Silicon present as soluble silicate is determined colorimetrically. The method has been described by Thompson and Houlton (1933) and by Wattenberg (1937). Because of the rapidity with which water samples are contaminated by silicate that dissolves from the glass, the analyses should be made soon after the water samples are collected. Waxed containers are sometimes recommended, and it is always desirable to use “aged” bottles that have been thoroughly leached with sea water. Tourky and Bangham (1936) tested the reaction between the molybdate reagent and colloidal silica and found that the color development was not proportional to the amount of silicon present. Treatment of the colloidal silica with alkali prior to analysis yielded correct values. Experiments with sea water indicated that colloidal silica may pass into true solution on ageing.

Fluorine is present in oceanic sea water in concentrations slightly above 1 mg/L. It is present as fluoride and, according to the work of Thompson and Taylor (1933), bears a constant ratio to the chlorinity.


The method of determination is described by these authors. Little is known concerning the role of fluorine in the sea.

Nitrogen occurs in sea water both in compounds of various kinds and as free dissolved nitrogen gas. As it is an essential constituent of living matter, nitrogen is found in organic compounds both in organisms and in particulate and dissolved organic material in amounts between 0.1 and 10.0 μg-atoms/L (p. 254). In addition, it is present as nitrate, nitrite, and ammonia. In routine observations only the inorganic nitrogen compounds are determined. Nitrate- and nitrite-nitrogen are determined colorimetrically, and the ammonia either colorimetrically (Robinson and Wirth, 1934) or by micro-titration after distillation (Krogh, 1934).

The nitrate method originally described by Harvey (1926) is given by Wattenberg (1937). Rakestraw (1936) and Wattenberg describe the procedure for the determination of nitrite. Since the inorganic nitrogen compounds are subject to change after the water samples have been collected, analyses must be run within a few hours. Even the addition of preservatives may not prevent changes in the NH3 and NO2, indicating that purely chemical transformations may be involved. Ammonia tends to disappear in storage, and nitrite sometimes decreases, but at other times shows an increase. The nitrate, which is more abundant, does not show such relatively large changes.

Because of their relatively low concentrations and their utilization by organisms, the inorganic nitrogen compounds show a wide range in values:

The distribution of nitrate in the oceans has been and is studied a great deal, as it may limit the production of phytoplankton when it is reduced to minimal quantities in the surface layers. Nitrate-nitrogen usually shows a subsurface maximum at a depth of several hundred meters. Nitrite nitrogen has a peculiar distribution and is generally found in a rather thin stratum in or above the thermocline. Lessi is known concerning the distribution of ammonia, as it is not so readly measured as the other inorganic compounds of nitrogen, but it is apparently rather uniform throughout the water column.

Nitrogen compounds are carried to the sea by rivers and by precipitation. The greater part of these are supposed to have been fixed by electrical discharges in the atmosphere. Possibly a certain amount of the fixed nitrogen in the sea is liberated as free nitrogen and returned to the atmosphere. Bottom sediments contain a small percentage of organic nitrogen in resistant organic detritus, and a part of this is

permanently lost from the water, as it is found in all types of sediments, both recent and fossil. As the carbon: nitrogen ratio in organic material is relatively constant, the organic nitrogen is frequently used as a measure of the amount of organic matter in marine sediments and also in the water. The distribution of nitrogen compounds and the nitrogen cycle in the sea are discussed in chapters VII and XVIII.

Aluminum is present in sea water in very small amounts. The colloidal clay particles which are carried to the sea contain a large percentage of aluminum, and hence analyses of water samples collected near shore may show the presence of aluminum, but it is not necessarily all in solution. The value given in table 36 is the average quantity reported by Haendler and Thompson (1939). Their values range between 0.006 and 0.065 mg-atoms/L (0.16 and 1.8 mg/L) with an average of 0.02 mg-atoms/L (0.54 mg/L).

Although earlier workers (Thompson and Robinson, 1932) were unable to detect rubidium in sea water, Goldschmidt (1937) has reported about 0.002 mg-atoms/L (0.2 mg/L).

Lithium content of sea water has been investigated by Thomas and Thompson (1933), who found 0.014 mg-atoms/L (0.1 mg/L).

Phosphorus, which is present in sea water as phosphate ions, is another of the essential constituents of living organisms, and its distribution in the sea is markedly affected by organic agencies. In addition to the nitrogen and silicon compounds, phosphate-phosphorus has been considered as one of the substances that may limit production of plant life. The inorganic phosphorus concentration varies from virtually zero at the surface, under certain conditions, to approximately 0.003 mg-atoms/L (0.090 mg/L) at subsurface levels when values are corrected for salt error. There is frequently a subsurface maximum similar to that in the distribution of nitrate-nitrogen. Phosphorus removed from the surface layers by phytoplankton is largely returned to solution on the death and decomposition of the organisms. It is supplied by rivers, and some is removed from the sea, as a small quantity is present in most marine sediments. In certain shallow areas, phosphatic concretions are found that contains a rather high concentration of phosphorus. The mode of origin of these concretions is not yet known. It has been suggested that in many regions the water is supersaturated in respect to tricalcium phosphate which, therefore, may be deposited inorganically (Dietz, Emery, and Shepard, 1942).

Phosphate phosphorous is determined colorimetrically. The method has been described by Robinson and Wirth (1935) and Wattenberg (1937). Cooper (1938a) has discussed the magnitude of the salt error. Phosphate analyses are frequently carried out as routine observations, and our knowledge of the distribution of phosphate in the ocean is fairly comprehensive. The rather scant knowledge we have concerning the

amount of phosphorus present as particulate or dissolved organic pbosphorus will be discussed in connection with the phosphorus cycle in the sea (chapter VII).

The amount of barium in sea water has been reported by Goldschmidt (1937) as 0.0004 mg-atoms/L (0.05 mg/L). This is lower than the values reported by Thompson and Robinson (1932). Barium occurs in marine organisms and it is a constituent of most marine sediments. In certain localities the deposits contain large amounts of barium sulphate in the form of concretions and nodules. The mode of formation of these structures is not yet understood.

The distribution and concentration of iodine in the sea has received a great deal of attention because of its important role in the physiology of man and terrestrial animals. Marine products are an important source of iodine-rich foods, The form in which iodine occurs in sea water is not yet clearly understood, but at least part of it is present as iodide and iodate. It is concentrated to a marked degree by marine plants, and for many years sea weeds have been used as a commercial source of iodine. The distribution and determination of iodine in sea water and marine organisms have been discussed by Closs (1931) and Reith (1930).

Arsenic content of sea water has been investigated by Rakestraw and Lutz (1933), who report values ranging from 0.15 to 0.3 μg-atoms/L (9 to 22 μg/L). This wide range is attributed to the fact that organisms may utilize arsenic in place of phosphorus. It is known to be a constituent of the tissues of many marine forms. The exact form in which arsenic occurs in sea water is not yet known.

Iron is an essential constituent of plants and has been considered as one of the substances that may limit the amount of plant production in the sea. Investigations show that at least part of the iron is not present in true solution, as it can be removed by ultrafiltration. Cooper (1937b) has pointed out that the amount of iron in true solution as ferric or ferrous salts is probably less than 2 μg/L, whereas the total iron present is generally about ten times this quantity. The amount present in the plankton may be as much as 16 per cent of the total iron of the water. Harvey (1937) considers that diatoms are able to adsorb and utilize colloidal iron. Iron is brought to the sea in relatively large quantities in the colloidal clay particles, and consequently considerable amounts of iron are found in the marine sediments. In many instances the iron content of the sediments is even higher than should be expected, indicating addition of iron through physical, chemical, or organic agencies. In inshore areas near the source of supply the total iron content of the water is sometimes much higher than that found in the open ocean. Methods for the determination of iron in sea water in its various forms have been described by Thompson and Bremner (1935a and b), Cooper (1935), and Rakestraw, Mahnke, and Beach (1936).


Manganese is apparently subject to concentration by marine organisms. Thompson and Wilson (1935) have reported values between 0.02 and 0.2 mg-atoms/L (1 and 10 mg/L). The value cited by Goldschmidt (1937) is 4 mg/L. Interest in manganese has been aroused by the occurrence of manganese nodules which are widely distributed in certain types of marine sediments, particularly in the Pacific Ocean.

The quantity of copper present in sea water probably lies between 0.02 and 0.2 μg-atoms/L (1 to 10 μg/L) (Marks, 1938, Wattenberg, 1938). Copper is an essential constituent of many marine organisms and it is also considered a factor in the life history of oysters, as a relatively high copper content of the water is apparently necessary for proper development of the larvae.

Much interest is attached to the content of radioactive elements in sea water, because deep-sea sediments are high in radium, compared to igneous rocks, and it is considered that the enrichment must be due to precipitation from the water of radium or its precursors. The radium content of sea water has been studied by many investigators, using various techniques, but it is only recently that methods have been sufficiently refined to yield trustworthy results. Studies by Evans, Kip, and Moberg (1938) and by Pettersson and Rona (Föyn et al, 1939) show that the radium content, measured by the radon emanation technique, varies between about 0.2 and 3.0 × 10−13 ‰ in sea water of salinity approximately 35 ‰. The low values are found in the surface layers, and it is suggested that organisms are responsible for a selective removal of this element. Both groups of workers found that organisms concentrate the radium about one hundredfold in their soft tissues. Calcareous structures show an increase in the radium:calcium ratio over that in the water. The maximum value listed above—namely 3.0 × 10−13 ‰— was found in water in contact with the sediments (Evans, Kip, and Moberg, 1938), and generally the radium content of the deeper waters is about 1 × 10−13 ‰.

Pettersson and co-workers (Föyn et al, 1939) have emphasized the importance of searching for the radioactive precursors of radium, as this element has the relatively short half-life period of only 1690 years. Of these elements uranium and ionium are probably the most; important, but thus far only uranium has been examined. Karlik (Föyn et al 1939) has analyzed a number of samples from various parts of the oceans and obtained for oceanic water a mean value of 1.5 × 10−6 ‰. Surface waters have a somewhat lower content than those from greater depths, but Karlik does not consider that the data are sufficiently adequate to show any differential removal. Studies of the dilute waters of the Baltic Sea showed that the uranium content was a function of the salinity.

Föyn and Rona (Föyn et al) have sought for thorium in sea water, but have been unable to detect it by the most refined methods. By examining

very large samples they have fixed an upper limit of 0.5 × 10−6 ‰ for this element. Older and apparently less accurate methods yielded considerably higher values.

The radium content of marine sediments and the theories concerning the deposition of radium and its precursors are discussed in the chapter on marine sedimentation.

Preparation of Artificial Sea Water

It is impossible to prepare solutions that exactly duplicate the properties of sea water because (1) the ions (salts) in which the elements occur in sea water are not always known, (2) elements that occur in sea water in small amounts are present as contaminants in other compounds in quantities which may far exceed those that should be added, and (3) many of the salts which must be added in fairly large amounts are hygroscopic or contain water of crystallization and are difficult to weigh accurately. The latter difficulty may be partially avoided by preparing concentrated solutions of these salts, determining their concentration by chemical analysis, and adding the required volume of the solution.

Although it would be of great interest to prepare solutions duplicating all the physical and chemical properties of sea water, it is generally not essential. In studies of certain of the physical-chemical properties, it is sufficient to add to the solution only the more abundant ions. In other instances—for example, when chemical methods are to be standardized—only one element or ion need be accurately known and other ions only approximately. Furthermore, in experiments with marine plants the major elements may not have to be closely controlled, but it will generally be necessary to know the concentrations of the biologically essential elements that are normally present in small amounts. If possible, natural sea water should always be used in physical or biological studies, but in the latter case it is sometimes desirable to enrich the water with certain of the plant nutrients (p. 235). Rogers (1938) has discussed various “modified ” types of solutions that are used in experiments on marine animals.

In table 37 are given three suggested formulae for preparing solutions approximating the composition of sea water. They have been adjusted to yield solutions of 19.00 ‰ chlorinity. The recipe of McClendon et al (1917), which has been used quite extensively, contains the nitrogen, phosphorus, and silicon needed by marine plants. Additional elements may be necessary but are probably always present as impurities. The formulae of Brujewicz (Subow, 1931) and of Lyman and FIeming (1940) contain only the major elements. The last-mentioned recipe corresponds to the composition of sea water given in table 35. The other formulae have not been adjusted to the composition presented in earlier sections of

this chapter. In all cases the reagents used should be examined for contaminants and, if necessary, purified.

McClendon et al (1917) Brujewicz (Subow, 1931) Lyman and Fleming (1940)
Salt g/kg Salt g/kg Salt g/kg
NaCl 26.726 NaCl 26.518 NaCl 23.476
MgCl2 2.260 MgCl2 2.447 MgCl2 4.981
MgSO4 3.248 MgSO4 3.305 Na2SO4 3.917
CaCl2 1.153 CaCl2 1.141 CaCl2 1.102
KCl 0.721 KCl 0.725 KCl 0.664
NaHCO3 0.198 NaHCO3 0.202 NaHCO3 0.192
NaBr 0.058 NaBr 0.083 KBr 0.096
H3BO3 0.058 H3BO3 0.026
Na2SiO3 0.0024 SrCl2 0.024
Na2Si4O9 0.0015 NaF 0.003
H3PO4 0.0002
Al2Cl6 0.013
NH3 0.002
LiNO3 0.0013
Total 34.4406 34.421 34.481
Water to 1,000.0000 Water to 1,000.000 Water to 1,000.000

Dissolved Gases in Sea Water

All of the atmospheric gases are found in solution in sea water. In addition to nitrogen and oxygen, the most abundant gases in the air, carbon dioxide is present in large quantities in sea water, chiefly combined as carbonates and bicarbonates. Of the rarer gases, ammonia, argon, helium, and neon have been reported in sea water, and hydrogen is undoubtedly present in minute quantities. In the absence of dissolved oxygen, hydrogen sulphide may be present, and it is possible that in stagnating water other products of putrefactive decomposition, such as methane, may occur.

Because of its importance in biological processes the dissolved oxygen distribution in the oceans has been examined intensively. Besides being an index to the biological history of the water, the general character of the distribution of oxygen in the deeper water is helpful in studies of currents and of mixing processes. The carbon dioxide distribution is of equal biological importance; its discussion begins on p. 192. Nitrogen has not been studied very widely, as it is apparently chemically inert. Argon is also inert, and is sometimes included with the nitrogen when the

dissolved gases are determined gasometrically. The presence of helium and neon has been confirmed by Rakestraw, Herrick, and Urry (1939).

Determination of Dissolved Gases. The content of dissolved oxygen is usually determined by the Winkler method, which depends upon the oxidation of manganous hydroxide by the dissolved oxygen. When acid is added, the oxidized manganese reacts with potassium iodide and sets free iodine, in amounts equivalent to the original dissolved oxygen content, which is determined by titration with sodium thiosulphate. The Winkler method is simple and extremely accurate if certain precautions are observed in handling the water samples and reagents (Thompson and Robinson, 1939).

Problems relating to the determination of carbon dioxide are discussed on p. 192.

Dissolved nitrogen cannot be determined by direct chemical methods, and hence gasometric techniques must be used. In general, the sea-water sample is acidified and all the gases are driven off by boiling or by applying a vacuum. The carbon dioxide is then absorbed in alkali, and the oxygen is absorbed in alkaline pyrogallol. The residual gas is sometimes considered as “atmospheric nitrogen,” although actually there are other gases, principally argon, mixed with it. Rakestraw and Emmel (1937) developed a method for determining the dissolved oxygen and nitrogen content of sea water by first extracting the gases and removing the carbon dioxide, then absorbing the oxygen on phosphorus and the nitrogen on molten lithium. The oxygen contents determined in this way agreed with direct Winkler analyses. The nitrogen determinations on saturated water samples showed results consistently lower than the saturation values according to Fox (1907); further studies (Rakestraw and Emmel, 1938b) indicate that Fox's tables are slightly in error. The gases remaining after the extraction of nitrogen are considered as “argon.”

The presence of hydrogen sulphide can be detected by its characteristic odor. A method for its determination has been described by Gaarder (1916). Although commonly referred to as hydrogen sulphide, a part, at least, will not be present as free gas but as sulphide or bisulphide of some base. A hydrogen sulphide system somewhat comparable to the carbon dioxide system must exist, but it has not yet been investigated.

The determination of ammonia is discussed in the section dealing with nitrogen compounds.

The units to be used in reporting the concentrations of dissolved gases are mg-atoms/L or (ml of gas at NTP)/L.

In some cases it is of interest to know the excess or deficiency of the concentration with respect to water of the same temperature and salinity in equilibrium with the normal dry atmosphere. The saturation values for oxygen and nitrogen are given in tables 38 and 39. If the saturation

values are known, the percentage saturation may be calculated. In certain problems it is desirable to know the partial pressures of the gases dissolved in a given water sample. The factors for computing these values are discussed on p. 190.

Chlorinity (‰) 15 16 17 18 19 20
Salinity (‰) 27.11 28.91 30.72 32.52 34.33 36.11
Temperature (°C)
–2 9.01 8.89 8.76 8.64 8.52 8.39
0 8.55 8.43 8.32 8.20 8.08 7.97
5 7.56 7.46 7.36 7.26 7.16 7.07
10 6.77 6.69 6.60 6.52 6.44 6.35
15 6.14 6.07 6.00 5.93 5.86 5.79
20 5.63 5.56 5.50 5.44 5.38 5.31
25 5.17 5.12 5.06 5.00 4.95 4.86
30 4.74 4.68 4.63 4.58 4.52 4.46
* mg-atoms of oxygen per liter = 0.08931 × ml/L.
Chlorinity (‰) 15 16 17 18 19 20 21
Salinity (‰) 27.11 28.91 30.72 32.52 34.33 36.11 37.94
Temperature (°C)
0 15.22 15.02 14.82 14.61 14.40 14.21 14.01
5 13.43 13.26 13.10 12.94 12.78 12.62 12.45
10 12.15 12.00 11.86 11.71 11.56 11.42 11.27
15 11.04 10.92 10.79 10.66 10.53 10.39 10.26
20 10.08 9.98 9.87 9.76 9.65 9.54 9.43
25 9.30 9.21 9.11 9.02 8.92 8.82 8.73
28 8.89 8.84 8.72 8.62 8.53 8.44 8.35
* mg-atoms of nitrogen per liter = 0.08929 × ml/L.

The dissolved oxygen in the sea varies between zero and 0.75 mg-atoms/L (about 8.5 ml/L), although in areas of low temperature and intense photosynthesis the content may exceed this upper limit. Nitrogen, which is apparently unaffected by biological processes, varies between 0.75 and 1.3 mg-atoms/L (8.4 and 14.5 ml/L). The total

carbon dioxide in oceanic waters varies between about 1.5 and 2.5 mg-atoms of C/L (34 and 56 ml/L). “Argon” varies between 0.2 and 0.4 ml/L, and the content of helium and neon in sea water is about 1.7 × 10−4 ml/L. The latter values apparently represent the saturation values. Hydrogen sulphide, which is present in the water under exceptional conditions, may occur in amounts of more than 1.0 mg-atom of S/L (22 ml/L) (Ström, 1936).

Factors Controlling the Distribution of Dissolved Gases. The following general factors control the distribution of dissolved gases in the oceans: (1) temperature and salinity, which determine the concentrations when the water is at the surface and in equilibrium with the atmosphere, (2) biological activity, which markedly affects the concentrations of oxygen and carbon dioxide, (3) currents and mixing processes, which tend to modify the effects of biological activity through mass movement and eddy diffusion.

Water in contact with the atmosphere will tend to reach equilibrium either by giving up or absorbing the individual gases until the water is just saturated. Although the zone of contact is a thin one, convective movements due to cooling, evaporation, or wind action may bring a layer of considerable thickness into equilibrium with the atmosphere. According to Henry's law the concentration, m, of a gas in a liquid is related to the partial pressure, p, of the gas and to the character of the gas and the liquid: m = csp. The numerical value of cs, the coefficient of saturation (absorption), depends upon the units for expressing the concentration of the gas in the solution and its pressure, and upon the chemical character of the gas and the temperature and salinity of the water.

Gas Percent of volume or pressure Partial pressure, Torr
Nitrogen 78.03 593.02
Oxygen 20.99 159.52
Argon 0.94 7.144
Carbon 0.03 0.228
Hydrogen, 0.01 0.088
100.00 760.000

With the exception of water vapor the relative composition of the atmosphere can be considered for practical purposes as constant (table 40). This does not strictly apply to carbon dioxide, relatively slight changes in the partial pressure of which have a pronounced effect upon

the amount in solution, and hence upon the hydrogen ion concentration and other properties (p. 202). Because of the variability in the water vapor pressure, the saturation is always assumed to take place from a dry atmosphere at standard pressure, namely, 760 Torr. The natural fluctuations and regional differences in the atmospheric pressure are neglected.

The solubilities of those gases, such as oxygen and nitrogen, which do not react chemically with the water or its dissolved salts decrease with increasing temperature and salinity. The solubilities of oxygen and nitrogen in sea water of different salinities over the normal range of temperature were investigated by Fox (1907, 1909). Fox's values for oxygen are still the accepted standards, but his data for nitrogen have been superseded by those of Rakestraw and Emmel (1938b). The solubility of carbon dioxide is greater than that of oxygen and nitrogen because it reacts with the water. Part of the carbon dioxide is present as free CO2 and H2CO3, but in sea water by far the greater part is present as carbonates and bicarbonate, and for the same partial pressure the total CO2 content of sea water is much greater than that of distilled water or neutral salt solutions. The content of free CO2 and H2CO3 decreases with increasing temperature and salinity. Argon is sometimes included with the “atmospheric nitrogen,” and, because its solubility differs from that of nitrogen, the values of the saturation coefficients will be slightly modified. Little is known concerning the other gases in sea water; however, both hydrogen sulphide and ammonia are very soluble gases and their saturation values can play no important part in their distribution.

In table 41 are given values of the saturation coefficients (absorption coefficients) for oxygen, nitrogen, and carbon dioxide in fresh and sea water at different temperatures. The values for oxygen are from Fox (1909), as are also the values for nitrogen in distilled water. The other nitrogen values are from Rakestraw and Emmel(1938b). The values for carbon dioxide (Buch et al, 1932) correspond to the total CO2 in water of zero alkalinity or to the free CO2 and H2CO3 in sea water. It is seen that carbon dioxide is much more soluble than the other two gases and that oxygen is about twice as soluble as nitrogen.

From table 41 it is seen that within the range of chlorinity normally encountered in the oceans the temperature is the most important property influencing the solubility (see also tables 38, 39).

In studies of the distribution of dissolved gases in the sea it is generally assumed that, whatever the location of a water particle, at some time it has been at the surface and in equilibrium with the air. In their studies of the dissolved nitrogen content Rakestraw and Emmel (1938a) have found that the water is virtually saturated (referred to a normal atmosphere), regardless of depth; therefore this assumption appears valid and also indicates that biological activity involving either fixation or production of nitrogen cannot be sufficient to affect significantly the concentration of this gas in the water. As the waters of the oceans appear to have been saturated with oxygen and carbon dioxide at some stage in their history when they were at the surface, the differences between the saturation values (computed from the temperatures and salinities) and the observed contents are measures of the changes which have been effected by biological agencies. The factors influencing the distribution of carbon dioxide are discussed in the following sections, and the distribution of dissolved oxygen will be considered in many places in the ensuing chapters.

COEFFICIENTS OF SATURATION OF ATMOSPHERIC GASES (cs) IN WATER (Concentrations of oxygen,[a] nitrogen[b], and carbon dioxide[c] as ml/L and mg-atoms/L in equilibrium with 760 Torr = 1 atmosphere of designated gas)
Temperature 12° 24°
Chlorinity (‰) O2 N2 CO2 O2 N2 CO2 O2 N2 CO2
ml/L mg-atoms O/L ml/L mg-atoms N/L ml/L mg-atoms C/L ml/L mg-atoms O/L ml/L mg-atoms N/L ml/L mg-atoms C/L ml/L mg-atoms O/L ml/L mg-atoms N/L ml/L mg-atoms C/L

aFox (1909).

bDistilled water, Fox (1909); sea water, Rakestraw and Emmel (1938b).

cBuch et al (1932), after Bohr. Concentrations represent amounts of free CO2 and H2CO3.

0 49.24 4.40 23.00 2.06 1715 77.0 36.75 3.28 17.80 1.59 1118 50.2 29.38 2.62 14.63 1.31 782 35.1
16 40.1 3.60 15.02 1.73 1489 66.8 30.6 2.75 11.56 1.33 980 44.0 24.8 2.22 9.36 1.08 695 31.2
20 38.0 3.40 14.21 1.64 1438 64.5 29.1 2.61 10.99 1.26 947 42.5 23.6 2.12 8.96 1.03 677 30.4

The Carbon Dioxide System


Although an extensive literature exists concerning the carbon dioxide system in sea water, publications prior to about 1929 are now chiefly of historic interest. The solution of the problems involved awaited not only the development of suitable analytical methods for the determination of the total carbon dioxide and the various forms in which it is present in sea water, but also the development of the theory and methods for studying the hydrogen ion concentration and certain general theories in physical chemistry. In the brief discussion to follow, only the salient features of the contemporary theories will be presented. These may be adequate for many purposes, but the investigations are not yet closed. Methods of analysis require further refinements, and in many cases fundamental constants must be more accurately determined.

Early investigators studying the carbon dioxide in sea water attempted to apply methods similar to those used for fresh water, where the carbon dioxide is largely present as free carbon dioxide that can be driven off by boiling, by applying a vacuum, or by bubbling through the water a stream of CO2-free gas. The use of such methods on sea water gave variable and conflicting results. It was later found that in order to drive off all the CO2 a strong acid must be added to the water, indicating that at least part of the carbon dioxide was present as the carbonate or bicarbonate of some basic cation. Methods were then developed for the determination of the total carbon dioxide and also for measuring the quantity present as carbonate and bicarbonate ions. It is now considered that the CO2 can exist in the following forms in sea water and that under any given set of conditions equilibria will prevail:

If the gases in sea water are driven off by some suitable method, the CO2 present as dissolved gas will be removed and the equilibria will be displaced until virtually all of the free CO2 and carbonic acid are removed and the bicarbonate is all converted to carbonate. If a strong acid is
added to sea water, the equilibria will be displaced toward the free CO2; consequently, if sufficient acid is added, all the CO2 is set free and can be determined either chemically or gasometrically. If an alkaline substance, such as sodium hydroxide, is added to sea water, the equilibria are shifted toward the carbonate, and the amount of carbonate ion will be increased. Although only a small fraction of the free CO2 is hydrated to form H2CO3, in the following discussion the free CO2 + H2CO3 will be referred to as carbonic acid and written H2CO3.

From the foregoing discussion it can be seen that the total CO2 in sea water does not follow Henry's law for the solution of gases in inert solutions. Nevertheless, the partial pressure of the carbon dioxide in sea water in contact with the atmosphere will tend to establish equilibrium with that in the air. If the pressure is increased the amount in solution will be greater, and if it is reduced the quantity of CO2 will decrease. The quantity present in a sample of water in equilibrium with a given carbon dioxide pressure will depend on the concentration of carbon dioxide bound base and the temperature and salinity of the water sample. If these factors are kept constant, the partial pressure of CO2 can be used as a measure of the total carbon dioxide content of the water.

Hydrogen Ion Concentration (pH) of Sea Water. Sea water is normally alkaline. Since both the H+ and OH ions play parts in the equilibria, any understanding of the carbon dioxide system requires knowledge of their concentrations. Pure distilled water dissociates into hydrogen and hydroxyl ions:

The ionic product [H+] × [OH], when the concentrations are expressed in chemical equivalents per liter, varies somewhat with temperature, but at 25°C is 10−14 (p. 198). In pure water or in any solution that contains equal concentrations of H+ and OH ions, the solution is said to be neutral. If the concentration of H+ is in excess of OH, the solution is acid, and if less it is alkaline. The ionic product is a known function of the temperature and salt concentrations; hence, if [H+] or [OH] is known, the other can readily be calculated. For expressing the hydrogen ion concentration, a logarithmic scale is commonly used, where pH is the logarithm of the reciprocal of the hydrogen ion concentration expressed as normality; that is, as equivalents per liter, formula. Thus, a neutral solution has a pH of approximately 7, an acid solution a pH less than 7, and an alkaline solution a pH greater than 7. It should be noted that a unit change in pH corresponds to a tenfold change in the hydrogen ion and hydroxyl ion concentrations.

The hydrogen ion concentration, or pH, of a solution may be determined in various ways, but all are essentially either electrometric or

colorimetric. The hydrogen electrode, which is the standard for measuring hydrogen ion concentration, cannot be used for sea water, as it involves bubbling gas through the solution and, thus, disturbance of the carbon dioxide equilibrium. Two other electrometric methods are available—the quinhydrone electrode and the glass electrode. The quinhydrone method is not accurate in the pH range normally encountered, and the glass electrode has not yet been extensively applied to the study of sea water (Ball and Stock, 1937, Buch and Nynäs, 1939). Hence, virtually all of our knowledge concerning the hydrogen ion concentration in sea water is based on colorimetric methods.

Certain organic compounds classed as indicators have the property of changing color over a given range of hydrogen ion concentration. So-called bicolor indicators have one color when in an “acid” solution and another color when in an “alkaline” solution. For any indicator the color change takes place over a definite range in pH, and, at the hydrogen ion concentration that is numerically equal to the dissociation constant of the indicator, equal quantities of both color phases are present in the solution. The range in pH for which various indicators can be used is described by Clark (1928), who also gives in detail the methods for preparing the indicators. For sea water, cresol red and phenol red are generally used, as they cover the pH range normally found in the sea. When working in high pH ranges, bromthymol blue is commonly used.

Three important properties of pH indicators must be known before they can be applied to sea water: the dissociation constant of the indicator, the effect of temperature upon this value, and the salt error. The presence of neutral ions in the solution has a pronounced effect upon the color and, hence, upon the apparent pH as determined by indicators. This is known as the salt error. In general, neutral salts increase the apparent dissociation constant of the indicator, and therefore give low pH readings. In practice a carefully controlled quantity of an indicator solution is added to a sample of sea water, and either the color developed is compared to a set of tubes containing the equivalent quantity of indicator in solutions of known pH or the sample is examined in a bicolorimeter. The accepted colorimetric technique used for determining the pH of sea-water samples and the correction to be applied for salt error and temperature effects are described by Buch (1937) and by Buch and Nynäs (1939). Because of the effects of changes in temperature and pressure upon the dissociation constants of carbonic acid (p. 200) the measured pH of a sea-water sample will differ from the pH in situ. The magnitude of the temperature correction is given by Buch (1937), and the effect of pressure upon the pH has been studied by Buch and Gripenberg (see Buch et al, 1932).

The pH encountered in the sea is between about 7.5 and 8.4. That is, the hydrogen ion concentration ranges from 32 × 10−9 to 4 × 10−9

equivalents per liter. The higher pH values are generally encountered at or near the surface. Where the water is in equilibrium with the CO2 in the atmosphere, the pH is between about 8.1 and 8.3, but higher values may occur when the photosynthetic activity of plants has reduced the content of CO2. Below the euphotic zone the pH shows a certain relationship to the amount of dissolved oxygen in the water. In regions where virtually all the oxygen has been consumed, and consequently where the total CO2 is high, as at depths of about 800 m in the eastern portions of the Equatorial and North Pacific, the pH approaches a minimum value of 7.5. This is a limiting value, because no more CO2 can be formed. Below the minimum oxygen layer there is generally a gradual increase in pH with depth. Under the peculiar conditions that may prevail in tide pools, bays, and estuaries, the pH sometimes exceeds the values cited above. Furthermore, in diluted water and in isolated basins where H2S is produced the pH may approach 7.0 or even fall in the acid range.

Alkalinity and Carbon Dioxide Components. The total amount of carbon dioxide in sea water, present either as free gas or bound, may be determined gasometrically after a strong acid has been added to the water to break up the carbonate compounds. Such a method has been described by Greenberg, Moberg, and Allen (1932). In order to determine the carbon dioxide components—namely, the amounts present as carbonic acid (including the free CO2), bicarbonate, and carbonate—titrations must be made. For a given sample of sea water the amount of a strong acid (usually HCl about 0.01 normal) necessary to reduce the pH to about 4.5 is independent of the total CO2. This amount of acid is required to set free the weak acids whose anions have been bound against basic cations. It is therefore not only a measure of the quantity of anions of weak acids in the sample, but also of the cations balanced against them. This quantity, when expressed as the number of milliequivalents of hydrogen ions (mg-atoms of H+) necessary to set free the ions of weak acids in a volume of water which at 20° has a volume of 1 L, is known as the alkalinity. This quantity has also been referred to as the titratable base, excess base, titration alkalinity, and buffer capacity. The term alkalinity has been adopted as the standard designation by the International Association of Physical Oceanography (1939). It should be noted that the term as here defined has no relation to the hydroxyl ion concentration or to the fact that sea water is normally alkaline.

A number of methods have been suggested for determining the alkalinity, and these have been summarised by Thompson and Robinson (1932) and Gripenberg (1937b). In general, they follow one of two techniques. Either the titration is carried out in the presence of the carbon dioxide, in which case the end point is taken at about 4.5, or the carbon dioxide is driven off. In the latter case a higher pH, about

7.0, is used. When the carbon dioxide is driven off, an excess of acid may be added, the solution boiled to free it of CO2, and then the excess acid determined by titration; or the measurement may be made directly upon the sample, which is held at the boiling point during the titration. Another method has been to add a known quantity of acid to the sample of sea water and then to determine the pH (Mitchell and Rakestraw, 1933). It has been implied that all of these methods will give the same value for the alkalinity, but this is not necessarily the case, and it is hoped that a standard method will be established.

The alkalinity bears a fairly constant relation to the chlorinity. The alkalinity: chlorosity factor for surface water has been determined by a number of workers and found to be close to 0.120 when the alkalinity is expressed in terms of milligram-atoms. The designation specific alkalinity that has been used in some cases is obtained by dividing the alkalinity, as mg-atoms/L, by the chlorinity, in g/kg, but such a mixed ratio should not be used. In water from greater depths the ratio may be somewhat higher than that given above, approaching an upper limit of 0.125 near the sea floor (Wattenberg, 1933). In brackish water the ratio may be increased tremendously if the river water is high in bound carbonate compounds. When the alkalinity: chlorosity factor is to be used as an index in studies of industrial pollution, the “normal” change in the ratio with concentration must first be established in samples of sea water diluted with unpolluted river water. Observations by Moberg and Revelle (1937) and by Wattenberg (1936) have shown that in oceanic water the increase in the Ca:Cl factor with depth is equivalent to the rise in the alkalinity:Cl factor. This indicates that changes in the alkalinity and calcium are of common origin—namely, precipitation or solution of CaCO3. Further material concerning the regional variations in the alkalinity: chlorosity factor is given on p. 208.

Because variations in the alkalinity: chlorosity factor in oceanic water are associated with corresponding changes in the calcium: chlorosity factor, the calcium content of the water may be computed from the alkalinity by the following expression:

This procedure for estimating calcium has been followed by Wattenberg (for example, Wattenberg, 1936).

It may be seen that the alkalinity: chlorosity factor of 0.1205 is the same as the HCO3:Cl factor given in table 35, when the bicarbonate is expressed as milligram-atoms of carbon per liter. The reason for this identity is that, in preparing the table, it was assumed that the pH of the water was such that only bicarbonate ions were present and, hence, would be equivalent to the alkalinity.


In discussions of the carbon dioxide system in sea water, the concentrations of the components have commonly been given in millimoles per liter. These are numerically identical with concentrations given as mg-atoms/L of carbon.

Salts of weak acids containing the following elements are known to occur in sea water: carbon, boron, phosphorus, arsenic, and silicon. Of these, salts of carbonic and boric acid only are present in sufficient concentrations to affect the magnitude of the alkalinity. For the present, we shall neglect the boric acid, which does not affect the alkalinity determination and which has to be considered in the carbonate system only at higher pH's. The alkalinity may then be taken as a measure of the concentration of bicarbonate and carbonate ions, and

where the brackets indicate molar concentrations—that is, gram-atoms of carbon per liter. The [H2CO3] may be determined by titration with sodium hydroxide and the [CO3] by titration with acid, using carefully controlled end points. The method is described by Greenberg, Moberg, and Allen (1932). It is shown later that the concentration of either H2CO3 or CO3 will be negligible when the other is present in significant quantities. Therefore we may write:
By substituting in these equations the measured quantities, the other components may be obtained.

Studies of the carbon dioxide system based on measurements made by these methods have been reported by Greenberg, Moberg, and Allen (1932) and Moberg, Greenberg, Revelle, and Allen (1934). In fig. 38 are shown vertical distribution curves for the carbon dioxide components, total carbon dioxide calculated from the titrations and by direct gasometric measurements, and the alkalinity at a station off the coast of southern California. The total CO2 obtained by the two methods agrees very well and shows a general increase with depth. In the upper layers there is an appreciable amount of CO3, but this decreases to zero at 200 m, and below this level H2CO3 occurs in quantities which increase with depth. The alkalinity is indicated for the upper 200 m, below which it corresponds, of course, to the curve showing the amount of HCO3. The increase of alkalinity with depth may be partly due to the biological precipitation of CaCO3 in the upper layers, but in this area it is principally associated with the increasing salinity.

We shall now proceed to a discussion of the laws governing the equilibria between the various carbon dioxide components, the alkalinity,

and the partial pressure of CO2. Thus far it has been assumed that the alkalinity was a measure of the equivalents of HCO3 and CO3 ions and of the cations bound against them, but this concept must be modified somewhat. Hydrogen and hydroxyl ions must also be taken into account, and, at the higher pH's, the boric acid as well. Since boric acid is a weak acid, only the first dissociation stage need be considered. The balance between the ions with which we are concerned may be written


Vertical distribution of the alkalinity and carbon dioxide components off southern California.

If we indicate the portion of the base (alkalinity) directly bound to the carbon dioxide components as ACO2, the relationship may be written (Buch, 1933a,b)

where all concentrations are in gram-atoms per liter. K′B is the apparent first dissociation constant of boric acid in sea water at the particular temperature and salinity prevailing, and KW is the ionic concentration product of water, [H+] × [OH], under similar conditions. According to Buch (1938) the ionic product in sea water at 20° can be computed from the following equation:
pKW decreases by about 0.035 for each l-degree rise in temperature
(Dorsey, 1940). The expression pKW bears the same relation to KW that pH does to [H+] and is the logarithm of the reciprocal of the ionic product. The same convention is used in expressing dissociation constants. The total concentration of boric acid [∑H3BO3] can be obtained from the chlorosity:
According to Buch the dissociation constant of boric acid can be expressed as
The Correction to be applied to the alkalinity to obtain ACO2, is appreciable at the higher pH's, as can be seen from fig. 39, where the quantities are given as milliequivalents per liter.


Concentrations of hydroxidebound (AOH) and borate-bound (AH2BO3) base as a function of pH in water of Cl = 19.00 ‰ at 20°C.

The equation for the first dissociation constant of carbonic acid is

and for the second:
where the brackets indicate molar concentrations. By introducing the relationships
it is possible to eliminate HCO3 and CO3 from the above equations and obtain them in the following form:
Extensive investigations have been carried out to determine the magnitudes of K1′ and K2′. These studies have been reported by Buch, Harvey, Wattenberg, and Gripenberg (1932), and by Moberg, Greenberg, Revelle, and Allen (1934). Buch and others have followed up the work,
using more refined methods and theories. According to Buch et al (1932),
Corrections for temperature can be computed from the following expressions:

The effect of hydrostatic pressure, as expressed by means of the depth Δz in meters, on the first dissociation constant is

Buch (1938) found that the second dissociation constant of carbonic acid in sea water at 20°C may be computed from the following equation:

Over the normal range of chlorinity in sea water a simpler expression is adequate—namely,

The temperature and pressure corrections are
For any given set of conditions, as expressed by means of temperature, salinity, pressure, alkalinity, and pH, the values of ACO2 and K′1 and K′2′ can be calculated. From these the total carbon dioxide and its various components can be computed from the following equations (Revelle, 1934):
In fig. 40 are shown the variations in the carbon dioxide components with pH, calculated from the equations given above for sea water of
Cl = 19 ‰ at ϑ = 20°C, and at atmospheric pressure. The CO2 components are given as mg-atoms/L of carbon and as ml/L of CO2.

The partial pressure of carbon dioxide is related to the amount of free CO2 + H2CO3 (indicated as H2CO3; see p. 190) in the solution:

The value of cs (p. 191) depends upon the temperature and salinity and the units used for expressing the concentration and partial pressure.


Carbon dioxide components in sea water of Cl = 19.00 ‰ at 20°C as a function of pH and the partial pressure of carbon dioxide.

In fig. 41 are shown curves for cs at different temperatures and chlorinities, where they represent the amount of H2CO3, in milligram atoms of carbon per liter of sea water, in solution under the designated conditions when the partial pressure of CO2 is 1 physical atmosphere (760 Torr). At 20° and 19 ‰ Cl, cs, is 34.2. That is, a partial pressure of one atmosphere of CO2 would be in equilibrium with a solution containing 34.2 milligram-atoms of carbon as free CO2 + H2CO3. The data are from Buch et al (1932).


The variations in pco2 with the other components is shown in fig. 40. The range is from less than 0.01 to greater than 100 Torr (0.1 × 10−4 to 1000 × 10−4 atm). The relationship over the pH range normally encountered in sea water is shown in the inset diagram in fig. 40. Between pH 7.5 and 8.3, pco2 decreases from 1.4 to 0.15 Torr (18.0 to 2.0 × 10−4 atm). The average partial pressure of CO2 in the air is about 0.23 Torr; hence surface sea water of Cl = 19.0 ‰ at 20° will have a pH of 8.2 if it is in equilibrium with the atmosphere.


Absorption coefficient (cs) of carbon dioxide in sea water as a function of temperature and chlorinity.

Sea water is a very favorable medium for the development of photosynthetic organisms. It not only contains an abundant supply of Co2, but removal or addition of considerable amounts results in no marked changes of the partial pressure of CO2 and the pH of the solution, both of which are properties of importance in the biological environment (p. 268). If the CO2 available for photosynthesis is assumed to be H2CO3 + ½HCO3, 0.48 mg-atoms of carbon per liter may be removed from water of Cl = 19 ‰ with an increase of the pH from 7.5 to only 8.5. In distilled water or in an inert salt solution of zero akalinity initially at pH 7.5, the total CO2 would be about one seventh of that amount.

Buffer Action of Sea Water. If a small quantity of a strong acid or base is added to pure water, there are tremendous changes in the numbers of H+ and OH ions present, but the changes are small if the acid or base is added to a solution containing a weak acid and its salts or a weak base and its salts. This repression of the change in pH is known as buffer action, and such solutions are called buffer solutions. Sea water contains carbonic and boric acids and their salts and is, therefore, a buffer solution. Let us consider only the carbonate system. Carbonate and bicarbonate salts of strong bases, such as occur in sea water, tend to hydrolyze, and there are always both H+ and OH ions in the solution. If an acid is added, carbonate is converted to bicarbonate and the bicarbonate to carbonic acid, but, as the latter is a weak acid (only slightly dissociated), relatively few additional hydrogen ions are set free. Similarly, if a strong base is added, the amount of carbonate increases, but the OH ions formed in the hydrolysis of the carbonate increase only slightly. The buffering effect is greatest when the hydrogen

ion concentration is equal to the dissociation constant of the weak acid or base—that is, when the concentration of the acid is equal to that of its salt.

Cycle of CO2 Between Sea and Atmosphere. Investigations of the partial pressure of CO2 in the ocean and the atmosphere have been made by Krogh (1904) and Buch (1939a,b). The following internal changes will increase or decrease the pCO2 in the surface layer:

Increase pco2 Decrease pco2
1. Rise in temperature 1. Decrease in temperature
2. Rise in salinity (evaporation) 2. Decrease in salinity
3. Respiration 3. Photosynthesis
4. Precipitation of CaCO3 4. Solution of CaCO3
5. Deep water brought to surface

The partial pressure of CO2 in the surface water can be computed with sufficient accuracy when the temperature, salinity, alkalinity, and pH are known, but, before a better understanding of the CO2 exchange between the sea and the atmosphere can be obtained, a far more comprehensive study of the partial pressure of the atmospheric CO2 must be made. Buch (1939b) has reported a number of direct observations on the CO2 content of the air which indicate that polar air is relatively low in CO2 (pCO2 = 0.23 Torr), compared to continental and tropical air (pCO2 = 0.25 Torr). It has been suggested that in low latitudes the air is enriched with CO2 from the ocean and that the general atmospheric circulation carries the CO2 into high latitudes. There it again dissolves in the sea water, which in time brings it back toward the Equator.

Activity of Ions in Sea Water. The apparent first and second dissociation constants of carbonic acid and the first dissociation constant of boric acid in sea water are larger than in distilled water and increase with increasing salinity. That is, the strength of these acids appears to be greater in solutions containing salts. These phenomena can be accounted for on the theory of activity introduced by Lewis and Randall (1923) and developed mathematically by Debye and Hückel. In a solution containing a mixture of electrolytes, such as sea water, there is a mutual interference of the ions, so that their activity or ability to participate independently in some reaction is much reduced. Most chemical determinations measure the total concentration of some ion and not its activity; however, certain physical measurements show the activity. For example, electromotive force determinations involve the activity of the hydrogen and other ions. Similarly, the measurement of the vapor pressure of a solution of a nonvolatile compound is an indication of the activity of the solvent. The activity will be less than that of the pure solvent under similar conditions. The partial pressure of dissolved gases is a measure of their activity.


The activity coefficient γ is related to the activity α of an ion as in the following example:

Hence, αH+ is the activity of the hydrogen ions expressed as gram-atoms per liter and [H+] is the stoichiometric concentration.

In the studies of the carbon dioxide system in sea water, the total CO2, alkalinity, and carbon dioxide components are measured chemically, and hence the values represent the stoichiometric values and not the activities. On the other hand, the hydrogen ion concentration is determined colorimetrically or electrometrically, and these methods yield the activity of hydrogen ions directly. Therefore, in the equations relating the CO2 components αH+ could have been inserted instead of [H+]. The dissociation constants have been referred to as apparent dissociation constants, in contrast to the thermodynamic constants that would be obtained at infinite dilution, where the activity coefficients (γ) are unity. The apparent dissociation constants are indicated by the symbol prime (′)—for example, K2. For carbonic acid the thermodynamic second dissociation constant can be written

and this is related to the apparent dissociation constant (Moberg et al, 1934):
The value of γ H2Co3 depends upon the relative solubility of CO2 in pure water and in sea water at the same temperature and upon the activity of the water in the two cases:
The subscript o indicates values at infinite dilution, and the subscript s the values at the concentration under consideration. The values of co and cs can be obtained from fig. 41, and es and eo, (the vapor pressures) can be computed (p. 67).

From the empirical equations relating pK′1 and pK′2 to the temperature and chlorinity, we know that at 20° and zero chlorinity the thermodynamic values are pK1 = 6.47 and pK2 = 10.288, and, at 19.00 ‰ Cl, pK'1 = 5.97 and pK'2 = 9.02. By substituting these values and γH2CO3

(in this case 1.131), it is possible to determine the activity Coefficients for the carbonate and bicarbonate ions. This substitution yields
Thus, in sea water of 19.00 ‰ Cl at 20°, only about one third of the bicarbonate and one fiftieth of the carbonate ions are “active.” This will be considered again with reference to the solubility of CaCO3.

Empirical equations have been presented which relate pK′1 and pK′2 to the cube root of the chlorinity. It has been shown (Buch et al, 1932, Moberg et al, 1934) that these equations are generally valid for salt solutions other than sea water if the ionic strength is used instead of chlorinity as a measure of concentration. The ionic strength (μ) of a solution is obtained by first multiplying the concentration of each individual type of ion, in moles per kilogram of solvent water, by the square of its valence, and then taking half the sum of these products (Lewis and Randall, 1923, p. 373). Lyman and Fleming (1940) have shown that the ionic strength of sea water in the normal range of concentration may be computed from the expression

Moberg et al (1934) have summarized the pertinent data, but no satisfactory expressions have yet been developed to show the manner in which the activity coefficients of the different ions in sea water change with concentration.

Solubility of CaCO3. The solubility of an electrolyte, such as calcium carbonate, may be expressed by a solubility product. The solubility product is identical with the ionic product (if concentrations are expressed as moles per liter) when the solution is in equilibrium with the solid salt and, therefore, saturated. The value of the solubility product depends upon temperature, the concentration of other ions (salinity), and the hydrostatic pressure. If, under a given set of conditions, the ionic product is less than the solubility product, the solution is undersaturated; if the ionic product is greater the solution is supersaturated, and if suitable nuclei are present, precipitation will proceed until the ionic product equals the solubility product.

The solubility product of CaCO3(KCaCO3) in distilled water at 20° is 5.0 × 10−9. In sea water of chlorinity 19.00 ‰ and at the same temperature, the calcium content is 10.23 mg-atoms/L, and at pH 8.2 the carbonate ion concentration is 0.26 mg-atoms/L of carbon. Therefore the ionic product is

which is 530 times greater than the solubility product in distilled water. If solubility data for distilled water are to be applied to sea water, it is
necessary to take the activity of the ions into account. It has been shown (p. 205) that the activity of the CO3 ions is about 0.02, and introducing this correction reduces the apparent supersaturation to about tenfold. Since no data are yet available for the activity of the Ca++ ions in sea water, it is impossible to apply solubility products determined for distilled water to test the relative saturation of calcium carbonate and other salts in sea water.

It is therefore necessary to determine empirically the concentrations of Ca++ and CO3 that can exist in contact with solid CaCO3. Working at 30° and increasing the CO3 content by lowering the total CO2, Revelle and Fleming (1934) obtained precipitation of CaCO3 as aragonite needles and spherulites. The calcium content of the solution was determined directly, and the CO3 was calculated from measurements of the pH, alkalinity, and chlorinity. The average of three experiments gave K'CaCO3 = 2.4 × 10−6 at 30°C. Wattenberg has carried out a number of studies on the solubility of calcium carbonate in sea water. His general procedure was to add CaCO3 crystals to sea water, and in some instances to increase the pCO2 The sea-water samples were then placed in sealed flasks and agitated until equilibrium was established. The calcium content of the water was calculated from the alkalinity measurements (p. 196), and the CO3 was obtained with the aid of pH determinations. Wattenberg and Timmerman (1936) give the following values for the apparent solubility product in sea water of Cl = 18.5 to 19.5 ‰:

Temperature, °C 0 5 10 15 20 25 30 35
KCaCO3 8.1 7.9 7.4 6.8 6.2 5.5 4.7 3.8 × 10−7

Ionic product [Ca++l × [CO3] in sea water of Cl = 19.00 ‰ at 20°C as a function of pH. The horizontal lines indicate the solubility product according to Wattenberg (0.62 × 10−6 and according to Revelle and Fleming (3.2 × 10−6).

At 30° Wattenberg's value is only one fifth of that obtained by Revelle and Fleming. No satisfactory explanation of this difference has yet been offered. In fig. 42 the ionic product in sea water of Cl = 19.00 ‰ at 20° is plotted against pH. The value of K'CaCO3, for this temperature obtained by Wattenberg—namely, 0.62 × l0−6—would indicate that supersaturation exists at all pH's above about 7.5. Surface water in equilibrium with the atmosphere has a pH of about 8.2 and would be

about sixfold supersaturated. The KCaCO3 obtained by Revelle and Fleming, on the other hand, if corrected to 20°, would be 3.2 × 10−6 which would indicate that surface water is about saturated with CaCO3 under these conditions. It hardly seems likely that surface sea water can be as highly supersaturated as Wattenberg's values would indicate. As an explanation it has been suggested that the calcium carbonate may not be free as ions but may be present in some complex, or possibly as colloidal calcium carbonate. On the other hand, it may be that the empirically determined values of K'CaCO3 are not applicable to sea water because of lack of equilibrium or faulty experimental technique. Further studies must be made before these points can be decided.

Smith (1940) has investigated the calcium carbonate deposition that takes place in the shallow waters overlying the Bahama Banks. He found the alkalinity to be much reduced, and from measurements he computed the ionic product, [Ca++] × [CO3]. Minimum values of the product that he considers approach the solubility product were found to fall between the values of Revelle and Fleming and those of Wattenberg.

According to Wattenberg (1936), K'CaCO3 at 20° changes with the chlorinity in the following way:

Cl, ‰ 0 5 10 15 20
KCaCO3 0.05 1.8 4.0 5.0 6.2 × 10−7

That is, in sea water the apparent solubility product of CaCO3 increases with chlorinity and decreases with temperature.

Revelle (1934) and Wattenberg (1936) consider that hydrostatic pressure has no significant effect upon the value of K'CaCO3. However, it should be remembered that, because of changes in the dissociation constants of carbonic acid, pressure does modify the relative amounts of HCO3 and CO3 in the water. Hence, water that is saturated at the surface will be slightly undersaturated if subjected to hydrostatic pressure, even if the total CO2 and calcium contents remain unaltered. Conversely, bottom water saturated with CaCO3 will be supersaturated when brought to the surface.

Although our knowledge of the values of K′CaCO3 under different conditions is incomplete and uncertain, it is possible to show the effect of changes in the conditions upon the ionic product, thus obtaining an understanding of those agencies that will favor precipitation or solution. Revelle (1934) has shown that, in surface water, increase of the salinity and the temperature, and decrease of the pco2—that is, the total CO2 content—all tend to increase the ionic product and therefore favor precipitation. The salinity effect is relatively small, and hence areas of high or rising temperature and of active photosynthesis will be those where precipitation of CaCO3 is most likely to occur. Opposite conditions will favor solution.


In the deep water the range in temperature and salinity is small, and therefore the variations in pCO2 will have the most pronounced effect. According to Wattenberg's studies (1933) the deep waters of the Atlantic Ocean are virtually saturated with calcium carbonate. Areas overlying red clay show slight undersaturation, and those overlying calcareous deposits (globigerina ooze) show either saturation or slight supersaturation.

Certain types of calcareous sedimentary material, both recent and fossil, do not show any evidence of organic origin. These are sometimes considered to be “chemical” deposits. In certain areas, microorganisms undoubtedly play an important part in establishing conditions that result in the incidental precipitation of carbonates. In tropical seas, this process may occur in shoal-water areas, coral reefs, lagoons, and mangrove swamps (Field, 1932). Although microorganisms can produce conditions favoring the precipitation of calcium carbonate, it is considered that they are effective agents in this process only in and on the sediments in the environments listed above. Smith (1940) has found that over the Great Bahama Bank CaCO3 is precipitated, under conditions of heating and excessive evaporation, on nuclei supplied by the sediments. The finely divided CaCO3 in certain deep-sea sediments is thought to arise from the break-down of the shells of foraminifera (p. 982) and not from precipitation in situ.

Distribution of Alkalinity, pH, and Carbon Dioxide Components. In the preceding discussion it has been assumed that the alkalinity: chlorosity factor is constant, and the value 0.1205 (milliequivalents per unit Cl) has been used in computations. This is in agreement with the average of a large number of observations made by Wattenberg (1933) in the Atlantic Ocean and those reported by Revelle (1936) for the upper layers of the Pacific Ocean. Wattenberg's value is commonly given as the specific alkalinity—namely, as milliequivalents per unit chlorinity, in which case it is 0.123. A number of different methods (p. 195) have been employed for measuring alkalinity, but there is no definite proof that they all yield similar values. Hence, it is difficult to compare the results obtained from various parts of the oceans by different workers. However, the findings of Wattenberg (1933), Mitchell and Rakestraw (1933), and Revelle (1936) all show that there is a somewhat higher alka1inity:chlorosity factor in the deeper water than there is in the surface layers. The increase in the factor usually amounts to about 0.005. Wattenberg examined a number of water samples collected immediately over the bottom, and in these he found an even larger factor. The low values in the surface waters are ascribed to removal of calcium carbonate by organisms possessing calcareous skeletons. Smith (1940) has suggested that, at least in certain localities, there may be inorganic precipitation if suitable nuclei are present when

the water is undergoing heating and evaporation. Whether or not calcium carbonate dissolves again while sinking through the water column depends upon the degree to which the water is saturated with this compound. The production of CO2, by metabolic activity at subsurface levels will favor solution. This production also occurs in the sediments and probably accounts for the sharp increase in alkalinity immediately over the bottom, which was detected by Wattenberg. Revelle (1936) has offered evidence which indicates that the alkalinity: chlorosity factor is higher in the North Pacific Ocean than it is in the Atlantic. The greater alkalinity may be attributed to the lower dissolved oxygen content of the intermediate and deep waters of the North Pacific. The lower oxygen content is indicative of a higher content of carbon dioxide, and hence greater solubility of calcium carbonate. Wattenberg's data indicate that the variation in the alkalinity: chlorosity factor with depth is much less in high latitudes, where organisms with calcareous skeletons are less abundant or are lacking and where vertical mixing produces more uniform conditions. That the change in alkalinity is analogous to variations in calcium content of the water has been pointed out by Moberg and Revelle (1937). River water contains relatively large proportions of calcium carbonate, and in regions of marked dilution there are great increases in the alkalinity: chlorosity factor. It may also be modified by the formation and melting of sea ice (p. 216).

The pH of sea water in contact with the air will vary between about 8.1 and 8.3, depending upon the temperature and salinity of the water and the partial pressure of carbon dioxide in the atmosphere. In areas of great dilution lower values may occur. At subsurface levels, where exchange of carbon dioxide with the atmosphere is impossible, the pH will vary with the extent to which the CO2 content of the water is modified by biological activity. In the euphotic zone, higher pH's are usually found; below this they decrease to a minimum corresponding in general to the layer of minimum oxygen content, and then increase again toward the bottom. Although variations in the salinity affect the pH, the predominant factor is the total carbon dioxide content or its partial pressure.

Unfortunately, differences in technique and in the constants used in arriving at pH's from colorimetric measurements make it difficult to compare the results of different workers. The most extensive studies made on the distribution of pH in the oceans are those of the Meteor (Wattenberg, 1933) in the Atlantic Ocean and those of the Carnegie in the Pacific Ocean. The distributions of pH and dissolved oxygen along a longitudinal profile in the eastern Atlantic Ocean are shown in fig. 43. The rather close similarity of the isolines can easily be seen, and similar patterns would be given by the partial pressure and total CO2 content, although in this case the relations would be inverse ones. The temperature

and salinity of the water are of secondary importance in determining the pH. In the North Pacific, where the oxygen content is low at intermediate and greater depths, when compared to the Atlantic, the pH is somewhat lower and approaches 7.5 where the oxygen content is reduced to less than a few tenths of a milliliter per liter. In stagnant basins (Ström, 1936) where large amounts of H2S are present, the pH may approach 7.0. The total CO2 content of water in contact with the air is chiefly dependent upon the alkalinity—that is, the salinity—and to some extent upon the temperature. Regional variations in the partial pressure of carbon dioxide in the atmosphere may cause some variations. Below the surface the total carbon dioxide content and, hence, the magnitude of the different components are primarily determined by the changes in the carbon dioxide content resulting from photosynthesis or respiration. If the temperature, salinity, and depth, and two of the following variables—COpH, alkalinity, or pCO2—are known, all components may be computed from equations given on p. 200.


Distribution of pH and dissolved oxygen in the eastern part of the Atlantic Ocean. The locations of the sections are shown in the map on the left-hand side. The vertical and horizontal scales are different in the two sections. (After Wattenberg.)

Solubility of Salts in Sea Water

The solubility of calcium carbonate in sea water has been examined in some detail, but relatively little is known about the other constituents. Because of the complex nature of sea water and the effect of other ions upon the activity of any one, the solubility product of a single salt in distilled water cannot be applied to sea water. Cooper (1937b) considers that most of the iron in sea water is not in true solution, but is present in some colloidal form, as the solubility product for the hydroxide is

extremely small. Wattenberg and Timmermann (1938) studied the solubilities of magnesium and strontium carbonates and magnesium hydroxide in sea water. The data in table 42 are from their work. It is interesting to note the great increase in the apparent solubility products in sea water, which is due to the reduced activity of the participating ions. For comparison, the ionic products for sea water (19.00 ‰ Cl, ϑ = 20°) at pH 8.2 have been computed. The apparent supersaturation of the calcium carbonate is discussed in the preceding pages. The ionic products of the other salts do not approach their solubility products. At pH higher than 9.0 the ionic product of Mg(OH)2 will exceed the solubility product, and hence removal of CO2 may result in precipitation of magnesium hydroxide as well as the carbonates.

Thompson and his co-workers (for example, Igelsrud and Thompson, 1936) have carried out extensive phase-rule studies of solutions containing some of the salts in sea water, but so far they have not extended their investigations to natural water.

Some indication of the great solubility of the major constituents is afforded by data on the separation of salts when sea water is frozen (p. 217). Somewhat similar data may be obtained from the evaporation studies by Usiglio (Thompson and Robinson, 1932), which again bring out the fact that sea water is far from saturated with most of the constituents.

Salt K Distilled water K' Sea water S = 35 ‰, ϑ = 20° Ionic product Cl = 19.0 ‰, ϑ = 20° pH = 8.2
CaCO3 0.5 × 10− 3 50 × 10− 3 270 × 10− 3
MgCO33H2O 0.1 × 10− 4 3.1 × 10− 4 0.14 × 10− 4
SrCO3 0.3 × 10− 9 500 × 10− 9 39 × 10− 9
Mg(OH)2 1 × 10− 11 5 × 10− 11 0.02 × 10− 11

The Oxidation-Reduction Potential of Sea Water

The oxidation-reduction potential is a measure of the ability of one chemical system to oxidize a second. It is generally expressed in volts relative to the normal hydrogen electrode. Those substances or solutions having high potentiaIs are able to oxidize those with lower potentials. Although considerable work has been done on the oxidation-reduction potentials in living organisms, little is known about the conditions prevailing in the water. The potential in sea water has been considered

by Cooper (1937a) to be associated only with the partial pressure of oxygen and the pH of the water. Under conditions of low oxygen content or when hydrogen sulphide is present, organic compounds dissolved in the water may have to be considered. The oxidizing or reducing conditions must be considered in two parts: namely, the intensity as expressed by the potential, and the capacity, or the poising of the system, which is a measure of the ability to oxidize or reduce a certain amount of material without significantly changing the potential. The poise of an oxidation-reduction system is somewhat analogous to the buffering capacity in hydrogen ion concentration. The oxidation-reduction potential is generally determined electrometrically, although in certain cases special colorimetric indicators can be used (Michaelis, 1930, Hewitt, 1937).

The oxidation-reduction potential of the environment is important to organisms. So-called aerobic bacteria thrive at a higher potential than micro-aerophiles, and anaerobic bacteria can exist only when the potential is low. Hence, in stagnant water and muds where there is no oxygen and the potential is low, only anaerobic forms can exist. The potential is also of geological importance, as the character of certain of the constituents of the sediments will be determined by the prevailing “oxidizing” or “reducing” conditions (p. 996).

Inorganic Agencies Affecting the Composition of Sea Water

The factors that may modify the absolute and the relative concen, trations of the substances in sea water are exchange with the atmosphere-inflow of river water, freezing and melting of sea ice, and biological activity. Biological processes and their effects upon the distribution of various elements are considered in chapter VII.

Exchange with the Atmosphere. The distribution of salinity in the oceans, and hence the concentrations of the major elements, is maintained by agencies that are described elsewhere, but one point must be considered at this time. Over the sea and along its shores, spray is continually being swept up into the air, and as the spray represents actual particles of sea water with its dissolved salts, this process affords a mechanism for the removal of salts from the sea. A large portion of the spray undoubtedly falls back into the water or is carried down by rain (Köhler, 1921). However, winds blowing toward the land will carry with them their content of salt, which may be deposited on the land directly or carried down by the rain. Observations by Jacobs (1937) on the chloride content of the air near the sea showed concentrations ranging between 0.07 and 0.5 mg of chloride per cubic meter of air. The amount increased with the wind velocity and was greatest with onshore winds.


A considerable proportion of the dissolved material carried to the sea by rivers is “cyclic salt”—that is, salt that has been carried inland by the atmosphere and then deposited or carried down by rain and snow (Clarke, 1924, Knopf, 1931).

Besides the exchange of salts that takes place between the atmosphere and the ocean as described above, there is an exchange of dissolved gases and nitrogen compounds which may modify the quantity of these substances present in sea water that is in contact with the atmosphere. The factors that affect the exchange of gases are described elsewhere. The exchange of water between the atmosphere and oceans was taken up in chapter IV.

Rain water contains relatively high concentrations of nitrogen compounds, which are believed to be formed from the constituents of the atmosphere by electrical discharges; hence the atmosphere supplies to the ocean, either directly through rainfall or indirectly through run-off from the land, a certain amount of fixed nitrogen. Whether this increment in the amount of fixed nitrogen is balanced by deposition of organic nitrogen in sediments or by the liberation of gaseous nitrogen through the decomposition of nitrogen compounds in the sea is not yet known.

Effects of Rivers on the Composition of Sea Water. The run-off from land is but a part of the cycle of leaching. The precipitation on the land contains only the cyclic salts, dissolved atmospheric gases, and nitrogen compounds. This water acts upon the rocks, contributing to the mechanical break-down of the solid material and extracting from them their more soluble constituents. The nature and quantity of the various elements dissolved depends upon the character of the rocks or soils with which the water comes in contact on its way to the sea. Because the leaching is carried out by water of low salt concentration yet relatively high in carbon dioxide compounds, it is capable of dissolving materials that would not pass into solution if they were in contact with sea water. In addition to dissolved material, rivers carry to the sea colloidal and particulate material in tremendous quantities. A considerable part of this debris is dropped to the sea bottom near shore, and much of the finer material coagulates and settles when mixed with sea water. Sea water reacts in various ways with the colloidal and finely dispersed material, and some of these reactions may affect the relative composition of the dissolved constituents. Interaction between the dissolved constituents of sea water and the sedimentary debris may be subdivided as follows: (1) solution of the constituents of the sediment, (2) adsorption on the sediment, (3) ionic exchange, and (4) reactions to form new substances. Little is known concerning the importance of these processes.

From the magnitude of the land area drained by rivers emptying into the sea and from the composition of the salts dissolved in river waters,

Clarke (1924) has estimated that rivers contribute to the sea each year 2.73 × l09 metric tons of dissolved solids. By weighting the composition of the water of different river systems in proportion to the total supply of dissolved material, he obtained the average composition of river water shown in table 43. Comparison with the composition of the dissolved solids in sea water shows, when referred to the halides, that there is an excess of all of the substances reported. Hence, the effect of river waters will be to modify the relative composition of the dissolved solids in sea water. Possibly a relatively large proportion of the carbon dioxide and nitrate should be regarded as cyclic. In addition, it should be recalled that the cyclic sea salts, carried into the atmosphere as spray and then deposited or carried down by precipitation onto the land, will enter the compilation. If we assume that all of the chloride in the river water is cyclic, the amounts of the other elements must be modified in the proportions in which they occur in sea water. It is hardly possible that this assumption is entirely correct, but it may yield values that are closer to the truth. The average composition of river water adjusted in this way for cyclic salts is included in the table.

Ion River water (weighted average) Sea water River water (less “cyclic” salts)
CO3 35.15 0.41 (HCO3) 35.13
SO4 12.14 7.68 11.35
Cl 5.68 55.04 0.00
NO3 0.90 ............. 0.90
Ca++ 20.39 1.15 20.27
Mg++ 3.41 3.69 3.03
Na+ 5.79 30.62 2.63
K+ 2.12 1.10 2.02
(Fe,Al)2O3 2.75 ............. 2.75
SiO2 11.67 ............. 11.67
Sr++, H3BO3, Br ...... 0.31 .....
100.00 100.00 89.75

It is not known whether the addition of dissolved solids brings about progressive changes in the relative composition of the sea salts or whether there is any progressive alteration of the total salt content or salinity. In any event, both processes must be exceedingly slow. The total amount of dissolved solids contributed by the rivers each year is only an infinitesimal fraction, 5.4 × 10−8, of the total dissolved solids in the ocean.

Even this small fraction does not represent a net gain, as there are many processes which remove material from solution. Certain elements accumulate in the marine sediments and are precipitated either through physical reactions or by biological activity. This is particularly true of the calcium and magnesium carbonate which form calcareous deposits and of the silicon which is found in organic siliceous deposits (radiolarian and diatom oozes). The sodium and potassium may be removed from solution by adsorption on, or ionic exchange with, the clay particles brought to the sea by rivers. However, in the latter case some other element will be released to the water in equivalent amounts. Other aspects of this problem are considered in connection with marine sediments and geochemistry, but it must be admitted that most questions related to changes in composition of sea water are still unsolved. In dealing with these it must also be considered that the amount of water may be changing. Essentially, all the river water is cyclic, but it is known that juvenile water of subterranean origin is continually being added to the surface water. In addition, the amount of water represented by the ice caps may be variable. Goldschmidt (1933) has estimated that for each square centimeter of the earth's surface there are 273 1 of water subdivided as follows:

Sea water … … … … … … … … … … … … … … … 268.45 1
Fresh water.… … … … … … … … … … … … … … 0.1
Continental ice.… … … … … … … … … … … … … 4.5
Water vapor.… … … … … … … … … … … … … … 0.003

The average composition of the river water is of interest in considering the effect on the oceans as a whole and over long periods, but particular investigations must be concerned with the effects brought about by individual rivers whose dissolved solids may differ markedly in composition and concentration from the average. Data can be obtained from Clarke (1924) or similar sources; as an illustration, values for several of the large American rivers are given in table 44.

From these examples it can be seen that the composition of individual rivers may differ considerably from the average. Thus, the Columbia River is low in chloride and the Colorado River is high, and the latter river is high in sodium and sulphate and below average in calcium and carbonate. The effect that dilution will have upon the chlorosity factors will therefore depend upon the character of the river water.

Thus far we have considered only the more abundant elements in the river water. Undoubtedly, all elements are carried to the sea either in solution or as finely divided particulate material. The high production of plant and animal life which frequently occurs near the mouths of rivers has sometimes been ascribed to the plant nutrients introduced by the rivers. Riley (1937) has found that the Mississippi

River water contains higher concentrations of phosphate and nitrogen compounds than the surface sea water in the Gulf of Mexico, and that this has a direct effect upon the amount of life in the waters off the mouth of the river. From studies in the English Channel, Atkins (1923) concluded, on the other hand, that the river water had little effect upon the immediate production, as most of the nutrients were used up by organisms living in the rivers.

Effects of Formation and Melting of Sea Ice on the Composition of Sea Water. A laboratory study of the freezing of sea water was made by Ringer, whose results have been reported by Krümmel (1907) and Johnstone (1928). In these experiments sea water was cooled in the laboratory, and at various temperatures below the initial freezing point the ice and precipitated salts were separated from the mother liquor. Sea water of salinity 35.0‰ begins to freeze at − 1.91°C (p. 66). At first, pure ice crystals separate, and, as the concentration of the brine is increased, the temperature must be further reduced to bring about the formation of additional ice. As the temperature is lowered and the concentration of the brine is increased, the solubility of certain of the dissolved salts is exceeded. At −8.2° the Na2SO4 begins to separate and continues to do so with further cooling. At −23° the NaCl begins to crystallize. In addition, a certain amount of CaCO3 precipitates. Ringer's analyses of the “ice” (including the ice crystals and the precipitated salts) and the brine when the temperature had been reduced to −30° are as follows:

Ion Average Mississippi River Columbia River Colorado River
CO3 35.15 34.98 36.15 13.02
SO4 12.14 15.37 13.52 28.61
Cl 5.68 6.21 2.82 19.92
NO3 0.90 1.60 0.49 ..........
Ca++ 20.39 20.50 17.87 10.35
Mg++ 3.41 5.38 4.38 3.14
Na+ 5.79 }8.33* 8.12 19.75
K+ 2.12 1.95 2.17
(Fe,Al)2O3 2.75 0.58 0.08 ..........
SiO2 11.67 7.05 14.62 3.04
Annual contribution of dissolved solids (metric tons) 100,000,000 19,000,000 13,416,000
Salt content (g/l) 0.166 0.0924 0.702
* Sum


One kilogram of sea water, initial salinity 35.05 ‰, yielded:

Ice crystals ……………………… 931.9 g
NeCl crystals … … … … … … … 20.23
Na2SO4 crystals… …………… 3.95
CaCO3 crystals.… … … … …… Trace
Brine… … … … ………………… 43.95

The brine contained 23.31 g of H2O and

Na+……………… 1.42 g Cl……………… 7.03 g
Mg++……………… 1.31 “ Br……………… 0.08 “
K+……………… 0.38 “ SO4……………… 0.03 “
Ca++……………… 0.39 “

From these data it is readily seen that, when the temperature of the ice and brine is lowered to −30°, there are marked differences in the relative composition of the salts in the “ice” and in the brine. If the cooling is continued to −50°, there is further separation of ice and salt crystals, but some very concentrated brine is still present.

From these experiments it would appear that the formation of sea ice might have a pronounced effect upon the relative composition of the salts in the water. The brine from which it formed would be modified in one direction, and, if melting took place in water other than that from which the ice formed, the effect would be in the opposite direction. However, the formation of sea ice in nature is not reproduced by these laboratory experiments. Let us suppose that, in a region where the depth to the bottom is moderate or great, sea water of normal composition is subjected to cooling at the surface. The resulting increase in density gives rise to convection movements that continue until the water at the surface reaches the freezing point, and then ice will begin to form. The brine, being of greater concentration but having virtually the same temperature, will sink and new water will be brought toward the surface and into contact with the ice. At first, isolated, elongated ice crystals are produced, but as the freezing continues these form a matrix in which a certain amount of the brine is mechanically included. The ice crystals themselves are at this stage probably “pure ice.” If the freezing proceeds rapidly, the brine will accumulate in separate cells within the body of the ice and, as the temperature of the ice near the surface is reduced, more ice crystals are formed, the cells decrease in size, and the concentration of the brine in the cells increases (fig. 16, p. 72). This may continue so far that solid salts crystallize in the cells. From this it can be seen that there is not necessarily any relative change in the composition of the dissolved salts in the sea water and in the sea ice (ice crystals plus the enclosed brine).

The salinity of the ice, using the same definition as applied to sea water, has been shown to depend upon the rate of freezing. Malmgren (1927), from the observations made by the Maud Expedition, gives the

following average values showing the relation between the salinity of new ice and the air temperature, where the latter is used aa a measure of the rate of freezing.

Air temperature (°C) Salinity of ice (‰)
− 16 5.64
− 28 8.01
− 30 8.77
−40 10.16

The salinities are based upon chlorinity determinations made on samples of the melted ice. The salinity of the surface water was about 30 ‰. The effect of the rate of freezing is also shown by the analyses of samples obtained in April from an ice floe that had started to form the preceding November:

Distance from surface of ice (cm) 0 6 13 26 45 82 95
Salinity of ice (‰) 6.74 5.28 5.31 3.84 4.37 3.48 3.17

The lower salt content of the deeper ice is related to the slower rate of formation. When ice is formed with extreme rapidity, its salinity will approach that of the water from which it is produced.

According to Ringer's experiments the cooling of ice containing cells of brine leads to formation of additional ice crystals and, if the temperature is reduced sufficiently, to separation of salt crystals within the ice. With very rapid freezing, brine and salt crystals may accumulate on the surface of the ice, making the surface “wet” at temperatures of −30° to −40°C and greatly increasing the friction against sled runners or skis.

In such rapidly frozen ice the cells containing the brine are large or numerous. If the temperature rises, the ice surrounding thg cells melts and the separated salt crystals are dissolved, but before complete solution has taken place the brine cells may join, permitting the brine to trickle through the ice. Under these conditions some of the solid salts may be left in the ice, and the composition of the water obtained by melting will differ from that of normal sea water. If, on the other hand, the temperature of the ice is raised to 0°C, all salts dissolve, the cells grow so large that the ice becomes porous, all brine trickles down from the portions of the ice above the sea surface, and the exposed old ice becomes fresh and can be used as a source of potable water.

Analyses by Wiese (1930) indicate that the processes that have been described may be effective in changing the relative composition of the salts, He found that the sulphate and alkalinity factors were greater in the ice than in the water and were greater in old ice than in newly frozen ice. This indicates that small amounts of sulphate, probably present as Na22SO4, have remained in the ice during the process of ageing and, probably, that the relative amounts of CaCO3 have changed.


Results of the Maud Expedition, reported by Malmgren (1927) and Sverdrup (1929), are not in agreement with the findings of Wiese. Chlorinities of water obtained by melting ice were systematically higher when determined by titration than when computed by means of Knudsen's Hydrographical Tables from observations of density. This discrepancy was interpreted to mean that the sea ice contains an excess of chlorides, but it may arise from the application of Knudsen's Tables to water that has been diluted by essentially distilled water, as explained on p. 59. The fact that the SO4/Cl ratio was nearly the same in the ice and the sea water (Malmgren, 1927, p. 9) supports the latter explanation and indicates that no changes in relative concentration had resulted from processes of freezing and melting. The problem, however, cannot be considered solved, and it offers opportunities for further laboratory investigations and observations in the field.

Geochemistry of the Ocean Waters

The total quantity of dissolved solids present in the waters of the oceans can be estimated by assuming an average salinity of 35 ‰ and assuming that the volume of the ocean is 1.37 × 109 km3 (p. 15). With a density in situ of 1.04 for the ocean waters, the dissolved solids amount to 5 × 1016 metric tons. This immense quantity of material would form a layer of dried salts 45 m thick over the entire earth, or 153 m thick over the present land area. The amount in tons of any element may be estimated by multiplying the value given in the first column in table 45 by 1.42 × 1012. The figures for the variable elements correspond to the higher values listed in table 36. Obviously the total amounts of even the trace elements are tremendous, and, if methods of extraction were economically feasible, the oceans would serve as an “inexhaustible” source of these substances.

Element Sea water S = 35‰ (mg/kg) Potential “supply” in 600 g of rock (mg/kg of sea water) Percentage in solution
Silicon 4 165,000 0.002
Aluminum 0.5 53,000 0.001
Iron 0.02 31,000 0.0001
Calcium 408 22,000 1.9
Sodium 10,769 17,000 65
Potassium 387 15,000 2.6
Magnesium 1,297 13,000 10
Titanium ............ 3,800 ?
Manganese 0.01 560 0.002
Phosphorus 0.1 470 0.02
Carbon 28 300 9
Sulphur 901 300 300
Chlorine 19,353 290 6700
Strontium 13 250 5
Barium 0.05 230 0.02
Rubidium 0.2 190 0.1
Fluorine 1.4 160 0.9
Chromium p 120 ?
Zirconium ............ 120 ?
Copper 0.01 60 0.02
Nickel 0.0001 60 0.0002
Vanadium 0.0003 60 0.0005
Tungsten ............ 41 ?
Lithium 0.1 39 0.2
Cerium 0.0004 26 0.002
Cobalt p 24 ?
Tin p 24 ?
Zinc 0.005 24 0.02
Yttrium 0.0003 19 0.002
Lanthanum 0.0003 11 0.003
Lead 0.004 10 0.04
Molybdenum 0.0005 9 0.005
Thorium <0.0005 6 0.01
Cesium 0.002 4 0.05
Arsenic 0.02 3 0.7
Scandium 0.00004 3 0.001
Bromine 66 3 2000
Boron 4.7 2 240
Uranium 0.015 2 0.8
Selenium 0.004 0.4 1
Cadmium p 0.3 ?
Mercury 0.00003 0.3 0.001
Iodine 0.05 0.2 25
Silver 0.0003 0.06 0.5
Gold 0.056 0.003 0.3
Radium 0.093 0.066 0.05
p = present

According to present theories, most of the solid material dissolved in the sea originated from the weathering of the crust of the earth. The problem as to the amount of rock weathered has been treated by Goldschmidt (1933) in the following way. For each square centimeter of the surface of the earth there are 278 kg of sea water; therefore, for each square centimeter the ocean water contains very nearly 3 kg of sodium. The average sodium content of igneous rocks is 2.83 per cent, and in sedimentary deposits it is 1.00 per cent. In the process of weathering, a certain amount of the material is leached away, and Goldschmidt estimates that the mass of the sedimentary deposits (Y) is 0.97 of the original igneous rocks (X) that gave rise to them. Therefore,

From this we find that for each square centimeter of the earth's surface about 160 kg of igneous rocks have been weathered. Therefore, approximately
600 g of rock have been weathered for each kilogram of water in the oceans. Of the total sodium, 65 per cent has accumulated in the sea water and 35 per cent has been deposited in sedimentary rocks. The 600 g of igneous rock have been, therefore, a potential supply of the constituent elements to the sea, although in most cases only a part of the material has actually dissolved or remained in solution. Using 600 g as the amount of rock weathered and following Goldschmidt's estimate (1937) of the composition of the earth's crust, the “supply” of elements listed in table 45 is obtained. A number of the minor constituents of rocks are not included in this tabulation. The “percentage in solution” has been obtained by dividing the amount of each element present in sea water by the potential supply. This procedure has been followed by Goldschmidt (1937).

Examination of table 45 shows that the elements may be grouped in three classes, depending upon the percentage in solution: (1) Sulphur, chlorine, bromine, and boron occur in amounts greater than those which could have been supplied by the weathering of the 600 g of rock. Goldschmidt considers that these elements were present in the primeval atmosphere as volatile compounds and that they accumulated in the ocean waters in the earliest times. (2) Calcium, sodium, potassium, magnesium, carbon, strontium, selenium, and iodine, which form relatively soluble compounds, are present in sea water in amounts greater than 1 per cent of the potential supply. (3) The remaining elements, which are present in small amounts.

It is striking that silicon, aluminum, and iron, the most abundant elements in igneous rocks (oxygen is actually the most abundant, but does not have to be considered here), are present in sea water in extremely small amounts. Thus, the relative abundance of the elements in sea water differs markedly from that in the earth's crust. With a few exceptions, all of the elements have been potentially available in much larger amounts than are actually present in solution. The relative composition of river water differs from that of sea water, and, in addition to the dissolved constituents, rivers introduce large quantities of particulate material that would pass into solution if the sea water were unsaturated with respect to these substances. Therefore, it appears that factors operating in the sea itself must control the concentrations of many of the elements that are potentially available in Iarge amounts. These factors are solubility, physical-chemical reactions, and biological activity. Our present knowledge is inadequate to designate which process or processes may control the concentration of a given element. Therefore, the following remarks will merely indicate the character of the factors that may be involved.

Certain elements may be present in such amounts that the solubility of their compounds may limit their concentration. In these eases,

additional amounts brought to the sea in solution by rivers will be removed by chemical precipitation. The quantities of other elements may be controlled by physical-chemical processes that are, however, more complex than the precipitation of some simple salt—for example, the reactions which may take place between the dissolved substances and the colloidal and particulate material introduced by rivers. Included among such processes are ionic adsorption, base exchange, and the formation of complex minerals. Such reactions may remove from solution ions that would not be precipitated in absence of colloidal or particulate material. Biological activity is undoubtedly of great importance in controlling the concentrations of many of the elements in the sea. Cyclical processes, in which elements are removed from solution but are later released by metabolic activity, need not be considered. However, a certain amount of the material built up by organisms falls to the sea bottom, becomes a permanent part of the deposits, and is therefore removed from solution. The concentration of elements carried down in this way may be considered to be at least partly controlled by the activity of marine organisms. The character of the skeletal structures and of the detrital organic matter deposited in this way is discussed in chapters VII and XX. Organisms remove from solution elements that would not otherwise precipitate, and, if conditions are such that some of this material becomes a permanent part of the sediments, it is obvious that biological activity must play an important part in controlling the composition of the water. Not only the major constituents of skeletal structures such as calcium, carbon, silicon, and so on, but nitrogen, phosphorus and many elements present in the sea in small concentrations are also accumulated by marine organisms.


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VII. Organisms and the Composition of Sea Water

Chemical Composition of Marine Organisms

Alterations of the concentration of the dissolved constituents of sea water are brought about by the development and subsequent death and disintegration of organisms. Virtually all of the substances extracted from the water are returned to solution by metabolic processes or by disintegration of the organisms, but the elements removed are returned to solution at some later time and often in some other part of the water column. Hence, the modifications may be in opposite directions at different times and localities. A small fraction of the organic remains accumulates on the sea bottom and is lost to the cycle.

Sea water probably contains in solution all of the chemical elements, although only some fifty have yet been detected. There is a large amount of data on the occurrence of various elements in marine plants and animals, but unfortunately the material is far from complete for any one biological group. Either a few elements only have been determined—such as iodine, for example, which has been thoroughly investigated—or only a portion of the organisms—for example, the skeletal structures—has been analyzed.

Vinogradov (1935, 1937) has compiled the chemical analyses of the lower plants and animals, both aquatic and terrestrial. He reports some sixty of the elements that have been found in one or more species. Webb and Fearon (1937) have tabulated thirty-nine elements that are commonly found and have divided these into two groups according to their apparent importance to living things: (1) eighteen invariable elements, and (2) twenty-one variable elements. These classes are further sub-divided on the basis of the concentration in which the elements are present. Seven elements are listed as contaminants (table 46).

The primary invariable elements are the essential constituents of carbohydrates, lipides (fats), and proteins. Some of the invariable elements classed as secondary or as microconstituents are always present in the lipides and proteins. This list is for plants and animals in general and not for marine forms alone. Comparison of tables 46 and 36 shows that nine elements (starred) detected in organisms have not yet been

reported for sea water, and that seven (uranium, thorium, cerium, lanthanum, yttrium, scandium, and radium), known to occur in sea water, are not listed by Webb and Fearon. Radium, at least, should be added to their list (p. 184).

Invariable (18) Variable (21) Contaminants
Primary 1–60% Secondary 0.05–1% Microconstituents <0.05% Secondary Microconstituents

* Not yet reported for sea water.

Hydrogen Sodium Boron Titanium[*] Lithium Helium
Carbon Magnesium Fluorine Vanadium Beryllium[*] Argon
Nitrogen Sulphur Silicon Zinc Aluminum Selenium
Oxygen Chlorine Manganese Bromine Chromium[*] Gold
Phosphorus Potassium Copper Cobalt[*] Mercury
Calcium Iodine Nickel Bismuth[*]
Iron Germanium[*] Thallium[*]

The lack of comparable data for the different types of organisms makes it necessary to consider their composition under three headings—namely, organic material (largely carbohydrates, lipides, and proteins), inorganic skeletal structures, and inorganic solutes in the body fluids. Although the proportions of carbohydrates, lipides, and proteins may vary considerably, the composition of any one type is rather constant, so that the average values in table 47 can be used with some confidence. Furthermore, there are numerous determinations of lipides (ether extract) and protein (based on nitrogen determinations), and from these measurements and the loss on ignition the carbohydrate may be computed. Skeletal structures differ so much in composition and in their mass, compared to that of the organic material, that they must be considered separately. Inorganic solutes in the body fluids are considered as a separate class, because they apparently do not differ very much in composition

and concentration from the surrounding sea water; consequently, the presence of water in the original sample will give merely rather high values for the inorganic solutes. The great changes in relative composition are found in the organic material and skeletal structures, for which reason, if a mixed plankton sample containing sea water is analysed, the results for the organic material and skeletal structures will not be materially affected if it is considered that all the sodium or chloride, as well as other elements in the proportions in which they occur in sea water, are in the body fluid or sea water.

Percentage composition Relative proportions by weight, C = 100
Element Carbohydrates Lipides Proteins Element Sea water Lipides Proteins
O 49.38 17.90 22.4 C 100 100 100
C 44.14 69.05 51.3 P 0.05 3.1 1.4
H 6.18 10.00 6.9 N 0.5 0.88 34.7
P 2.13 0.7 S 3150 0.45 1.6
N 0.61 17.8 Fe 0.07 0.2
S 0.31 0.8
Fe 0.1

In table 47 are given the average compositions of the three great classes of organic material (Rogers, 1938) and the relative proportions in which their component elements occur in sea water. The oxygen and hydrogen are not considered, and the values are adjusted to C = 100. The values for C, S, and Fe are from table 36; those for N and P are the average winter values in the English Channel (pp. 252, 258). In the lipides, phosphorus is concentrated, and in the proteins the nitrogen and phosphorus show a great increase with respect to carbon. The fact that sulphur, which is one of the relatively abundant elements in sea water, is a minor constituent of the lipides and proteins in organic material indicates that carbon, here used as the reference element, is itself markedly concentrated. The values given in table 47 are general averages, and those for marine organisms may differ slightly. It should be noted that changes in the proportions of carbohydrates, lipides, and proteins will modify the ratios in which the above-mentioned elements will be removed from the water. Many of the other elements that are concentrated by organisms—for example, iodine, iron, and copper—probably form a part of the organic material or they occur in the skeletal structures, as it is difficult to see how the free ions could be retained in the body fluids

at such tremendously higher concentrations than those in the surrounding sea water.

In table 48 are given analyses of certain types of skeletal material. In each case there is some organic matter, which is highest in the lobster carapace, and even in the phosphatic brachiopod shell it forms a large fraction. Of course, there are wide ranges in the proportion of inorganic skeletal structures in the whole organism, and in some cases such structures may be entirely lacking. The first three examples are for calcareous types with CaCO3 predominating, but in some groups MgCO3 forms an important part of the shell. The lobster may be considered as representative of the arthropods in general, although the proportion of organic matter is probably even greater in the small forms. The amount of phosphate is notable in the lobster and even more so in the phosphatic brachiopod shell, which is predominantly calcium phosphate. The sponge spicules are virtually pure hydrated silica and may be taken as representative of the diatom and radiolarian skeletons. The silica, iron, and aluminum in the other analyses probably represent impurities introduced by the presence of clay and sand grains. These analyses cannot be regarded as complete, and further examination will undoubtedly reveal many other elements present in small amounts. It should be noted that chlorine and sodium, the two most abundant elements in sea water, are not shown in any of these analyses. These elements form soluble compounds and hence would not be suitable for skeletal structures. From table 48 it can be seen that the development or re-solution of skeletal structures of marine organisms may be expected to affect the concentrations

of calcium, magnesium, carbon (as carbonate), sulphur, phosphorus, and silicon. Because of their relatively greater abundance in sea water, the distribution of magnesium and sulphur cannot be expected to show any appreciable effects of biological activity, but the distributions of the other elements mentioned do show certain features that can be attributed, in part at least, to the fact that they are important constituents of skeletal structures.

PERCENTAGE COMPOSITION OF SKELETAL MATERIAL (Recalculated from Clarke and Wheeler, 1922)
Substance Foraminifera (Orbitolites marginatis) Coral (Oculina diffusa) Calcareous alga (Lithophyllum antillarum) Lobster (Homarus sp.) Phosphatic brachiopod (Discinisca lamellosa) Siliceous sponge (Euplectella speciosa)
Ca 34.90 38.50 31.00 16.80 26.18 0.16
Mg 2.97 0.11 4.36 1.08 1.45 0.00
CO3 59.70 58.00 62.50 22.40 7.31 0.24
SO4 0.68 0.52 4.43 0.00
PO4 tr tr tr 5.45 34.55 0.00
SiO2 0.03 0.07 0.04 0.30 0.64 88.56
(Al,Fe)2O3 0.13 0.05 0.10 0.44 0.32
Organic matter, etc 2.27 3.27 1.32 53.45 25.00 10.72

The relative concentrations of the elements that are abundant in the body fluids do not differ very much from those in sea water (table 49). Although little is known concerning the less abundant elements, it appears that the inorganic portion of the body fluids can be considered as slightly altered sea water. Therefore, this part of the organism cannot play any appreciable part in modifying the composition of the water. Although the composition and concentration of the inorganic solutes is of no particular importance in the present problem, these features have been studied intensively in problems of osmotic pressure relations (chapter VIII) and in connection with the mechanism of solute and water exchange between aquatic organisms and their environment. These fields have been reviewed by Rogers (1938).

RELATIVE COMPOSITION OF BODY FLUIDS (Adjusted to Na = 100. Data from Robertson, 1939)
Element Sea water Echinus esculentus (Sea urchin) Homarus vulgaris (Lobster) Cancer pagurus (Crab)
Cl 180 182 156 156
Na 100 100 100 100
Mg 12.1 12.0 1.5 5.7
S in SO4 8.4 8.5 2.2 6.7
Ca 3.8 3.9 5.0 4.8
K 3.6 3.7 4.7 4.0

Thus far only the various fractions of the organisms have been discussed, and it is of interest to consider the composition of the entire plant or animal. As the plants are the primary “consumers” of inorganic material, it would be desirable to know the composition of such important groups as the diatoms and peridinians, but no complete analyses of these forms have been made. What information we have will be discussed below. The data for animals is also far from complete, but in table 50 are given three examples. The relative compositions have been adjusted to Na = 100, and, for comparison, the constituents of sea water are given in the same way. The relatively high proportions of the elements abundant in sea water which occur in the copepod analysis indicate the presence of considerable sea water in the original sample. As Archidoris possesses internal calcareous structures, the calcium content is high. Consequently the proportions of the elements constituting the organic material are rather low in these two cases. It is immediately obvious, however, that the essential constituents of the organic material, such as carbon, nitrogen, and phosphorus, are very high when compared to their relative concentrations in sea water.

Element Calanus (Copepod) Vinogradov, 1938 Fish[a] (Average) Archidoris britannica (Nudibranch) McCance & Masters, 1937–38 Sea water Concentration factors
Copepod Fish Nudibranch

aFrom analysis of Jowett and Davies (1938), Clements and Hutchinson (1939).

bWinter values, English Channel.

cHigher values from table 36.

Cl 194 180 180 1.1 1.0
Na 100 100 100 100 1.0 1.0 1.0
Mg 5.6 36 156 12.1 0.46 3.0 12.9
S 25.9 259 7.1 8.4 3.1 31 0.85
Ca 7.4 52 262 3.8 1.9 13.7 69
K 53.7 383 20 3.6 15 109 5.5
Br 1.7 0.6 3
C 1113 ca 4100 ca 480 0.26 4,300 15,800 1,850
Sr 11 0.12 92
Si 1.3 0.001[b] 13,000
F 69 0.01 6,900
N 280 1276 107 0.001[b] 280,000 1,276,000 107,000
P 24.1 256 6 0.0001[b] 241,000 2,560,000 60,000
I 0.04 0.0005 80
Fe 1.3 1.3 0.23 0.0002[c] 6,000 6,000 1,000
Mn 0.0008 0.0001[c] 8
Cu 0.008 0.43 0.0001[c] 80 4,300


If the relative amounts of the various elements in these animals are divided by their relative concentrations in sea water, a series of concentration factors, referred to sodium, are obtained. It will be seen that chlorine would give virtually identical results. The concentration factors range from about unity up to over two million for phosphorus in the fish, and in all three cases they are greatest for nitrogen and phosphorus. If it is assumed that the rates of diffusion of all substances and their rates of absorption by the organisms depend only upon the amounts of the ions in the water, then the concentration factors should be a measure of the time required to accumulate them. Those elements having the highest concentration factors would then be the ones that might limit the rate of growth. The data in table 50 indicate that nitrogen and phosphorus may very well be limiting elements in the sea, although it should be remembered that the examples are for animals that must obtain their supply of these elements either directly or indirectly from the plants. If the total carbon in sea water had been used as the reference element, only nitrogen and phosphorus would have significantly larger factors. But it is obvious that carbon is itself concentrated more than one thousandfold with reference to the major elements in sea water. According to table 50 the relative concentration

in sea water of the following elements may be markedly affected by biological activity: carbon, silicon, fluorine, nitrogen, phosphorus, iron, and copper. Other elements might be included if the analyses were more complete or if other types of organisms were examined.

Element Sea water Diatoms Peridinians Copepods Concentration factors (referred to carbon)
Diatoms Peridinians Copepods

aWinter values, English Channel.

bHigher value, table 36.

C 100 100 100 100 1 1 1
N 0.5[a] 18.2 13.8 25.0 36 28 50
P 0.05[a] 2.7 1.7 2.2 54 34 44
Fe 0.07[b] 9.6 3.4 0.13 137 49 2
Ca 1420 12.5 2.7 0.66 0.01 0.002 0.0005
Si 0.4[a] 93.0 6.6 0.13 232 16 0.3

In table 51 are given the relative concentrations of certain elements in diatoms, peridinians, and copepods, adjusted to C = 100. The data for the photosynthetic forms are recomputed from Vinogradov (1935), and for copepods are the same as those given in table 50. The concentration factors for nitrogen and phosphorus are about the same in all three forms. In the diatoms, iron is higher, while silicon has the highest factor, which may indicate that these elements also limit the rate of growth. For the peridinians the factors for nitrogen, phosphorus, and iron are nearly the same.

Interrelations Between Elements Whose Distribution Is Affected by Biological Activity

Because the relative composition of organisms living in the sea differs from that of sea water, their growth will tend to modify the composition of the water. The ultimate regeneration of the inorganic substances by biological processes will return the elements to solution, but the net effects will usually be in opposite directions at different times and in different parts of the water column. Tables 50 and 51 show that certain elements present in the water in low concentrations, such as nitrogen, phosphorus, iron, and silicon, are those removed in the largest relative amounts. The distribution of these elements, known as the plant nutrients, is profoundly affected by biological activity, their concentrations are virtually independent of salinity, and they are commonly referred to as nonconservative, in contrast to those elements that bear a constant ratio to the total dissolved solids.

Plants are the most important “consumers” of the inorganic substances. Their activity is restricted to the upper layers of the sea (the euphotic zone), where there is adequate light for them to carry on photosynthesis. In nearshore areas the thickness of the euphotic layer may be only a few meters, and even in the open sea, where the transparency is great, the growth of plants is restricted to the upper few hundred meters (chapter XVI). Animals living below the euphotic layer may remove elements from solution which are necessary for the secretion of skeletal structures, but most of the materials must come directly or indirectly from plants that develop near the surface. The metabolic activities of the plants, animals, and bacteria return the elements to inorganic form. Part of the regeneration must occur in the euphotic layer, but there is a general downward movement of the particulate matter, either living or dead, and, consequently, a continuous transport of the elements away from the surface layer. As described thus far, it

would appear that there is a continuous drain upon the resources of the surface layers without any mechanism for renewing the store of essential elements in the euphotic zone. Precipitation and rivers contribute a certain amount of these elements, but this is negligible in comparison to the quantities that are brought up to the surface layers by processes of vertical diffusion, convection overturn, and upwelling. In regions where these processes are active, there is a plentiful supply of nutrients; consequently, such areas can support very large populations and are said to be very productive, in contrast to the open oceans, where there may be only a meager supply of nutrients.

Redfield (1934), following an earlier suggestion by Harvey, showed that regardless of the absolute concentrations a constant ratio exists between the nitrate-nitrogen and phosphate-phosphorus content of sea water, that these elements are apparently removed from the water by organisms in the same proportions in which they occur, and that on the death and decomposition of the organisms they are returned to solution simultaneously. Cooper (1938a) proposed a modified ratio, pointing out that the phosphorus data for sea water used by Redfield had not been corrected for salt error. Fleming (1940) obtained from examination of additional data a slightly different relationship for the N:P in plankton. All these figures are given in table 52, which also shows the relation of carbon to the other two elements in plankton.

Source By weight By atoms
Redfield (1934) Plankton 53.2 8.2 1
Redfield (1934) Seawater 9.0 1 20 1
Cooper (1938a) Seawater 6.8 1 15 1
Fleming (1940) Phytoplankton 42 7 1
Fleming (1940) Zooplankton 40 7.4 1
Fleming—Average: Plankton. 41 7.2 1 106 16 1

The ratios given above hold very well for the nitrate and phosphate in ocean waters (see fig. 51), but, since they represent the net effect of biological activity, marked deviations from the ratios may be found in individual types of organisms. However, they indicate the order of magnitude of the relationships in marine organisms.

In order to extend the usefulness of these relationships, it is worth while to add the oxygen. As an approximation, it may be assumed that two atoms of oxygen are required for the oxidation of each atom of

carbon, and that in photosynthesis the same amount is released for each atom of carbon converted to organic matter. Gilson (1937) has pointed out that the amount of oxygen should be increased about 25 per cent to allow for the oxidation and reduction of nitrogen, but we shall not introduce this factor. Therefore,

In the euphotic layer, for each milligram of phosphorus utilized in photosynthesis, these ratios indicate that the plants will take up 7.2 mg of nitrogen (chiefly nitrate) and 76 ml of CO2 and release the same volume of oxygen. At lower levels, where regeneration is taking place, the consumption of 76 ml of O2 should set free the corresponding amounts of CO2, N, and P. The oxygen saturation value for water of 5° temperature is of the order of 7.0 ml/L; hence it may be seen that, if all of the oxygen has been consumed in subsurface water of approximately this temperature, the NO3-N and PO4-P will be increased by about 50 μg-atoms/L (0.650 mg/L) and 3 μg-atoms/L (0.090 mg/L), respectively. These values are approximately the largest amounts ever encountered in the ocean. If all the waters leaving the surface were saturated with oxygen and completely depleted of nitrate and phosphate, it might be expected that there would be a close agreement in the deeper waters between these substances and the oxygen depletion (difference between the saturation value and the observed content). Such a general relationship exists in waters that have left the surface in lower latitudes, but in higher latitudes water sinking from the surface is saturated with oxygen and contains appreciable amounts of nutrients; hence the relationship between the oxygen depletion and the nutrient content must have the form

where V is a variable amount in the water that sinks from the surface. In those regions in which there is a well-defined layer of minimum oxygen content at subsurface levels, layers of maximal nitrate and phosphate usually exist at or somewhat below the oxygen minimum.

The above comments do not necessarily apply to elements composing hard “inorganic” skeletal structures. Both calcium carbonate and silica are utilized by organisms in the euphotic layer and elsewhere, but the ratios of utilization of Ca, C as CO3, and Si with reference to, say, phosphate-P, depend upon the character of the organisms. As pointed out elsewhere (p. 208), CaCO3 is removed from the surface layers, and the same is true of the SiO2. Although there is generally a depletion of Si in regions where the nitrate-N and phosphate-P are low, the processes of re-solution of calcareous and siliceous structures do not necessarily parallel decomposition and the regeneration of the elements found in the soft parts of the organisms. Therefore, the general distribution of silicon in the sea differs somewhat from that of phosphate and nitrate, and the ratios between Si and N and Si and P are variable.


Locations of vertical sections and stations used to illustrate the distribution of phosphate, nitrate, and silicate, in the oceans.


Distribution of Phosphate, Nitrogen Compounds, and Silicate in the Oceans

The distribution of phosphate in the three oceans may best be shown by means of longitudinal vertical sections whose locations are indicated in fig. 44. The representations are intended to bring out only the major features of the vertical distribution, and for this reason many of the minor irregularities have been omitted. The section in the Atlantic Ocean (fig. 45) is based on data obtained by the Discovery (Deacon, 1933) in the Southern Hemisphere, by the Atlantis (Seiwell, 1935) in the North Atlantic, and by the Meteor (Defant et al, 1936) in the area to the south of Greenland. The section in the Pacific Ocean (fig. 46) has been constructed from Discovery observations in the Antarctic (Clowes, 1938) and from those of the Carnegie (in press). The section in the Indian Ocean (fig. 47) is based on Discovery observations (Clowes, 1938). Examination of these sections and the vertical distribution curves in figs. 48 and 50 shows that in general the distribution of phosphate and nitrate is characterized by four different layers: (1) a surface layer in which the concentration is low and relatively uniform with depth, (2) a layer in which the concentration increases rather rapidly with depth, (3) a layer of maximum concentration that is usually located somewhere between 500 and 1500 m, and (4) a thick bottom layer in which there is relatively little change with depth. Examination of figs. 45, 46, and 47 shows that the surface layer is thickest in mid-latitudes

in both hemispheres, that in these regions the layer of increasing concentration is clearly defined, and that the gradient is large. Associated with the divergences at and near the Equator, the surface layer is thin and the underlying gradient is very steep. This feature is brought out more clearly in fig. 198 (p. 710), which shows the vertical distribution of properties in the upper 300 m in a section across the equatorial currents in the Pacific Ocean. In high latitudes (above about 50°), where the surface layer of low concentration and the layer of rapid increase may be entirely lacking, high values of phosphate are found at the surface.


Phosphate distribution in a longitudinal section in the central Atlantic Ocean. Units: μg-atoms of phosphorus per 20° liter.


Phosphate distribution in a longitudinal section in the Pacific Ocean. Units: μg-atoms of phosphorus per 20° liter.


Phosphate distribution in a longitudinal section in the western Indian Ocean. Units: μg-atoms of phosphorus per 20° liter.

In the Atlantic Ocean the highest phosphate content (about 2 μg-atoms/L) is found in an intermediate layer extending northward from the Antarctic and centered at depths of about 1000 m. At all depths of about 1000 m or more there is a gradual decrease in phosphate concentration from south to north. A layer of minimum concentration containing less than 1.0 μg-atoms/L of phosphorus extends southward beneath the layer of maximum content. The phosphate concentration in the Antarctic is about twice that found in the northern part of the ocean. Regional differences in the phosphate distribution in the upper 50 m of the Atlantic Ocean are shown in fig. 217 (p. 787). Such differences are of importance in the distribution of plankton. The variations in a transverse section across the South Atlantic are illustrated in fig. 218 (p. 788), which shows the variable thickness of the surface layer of low phosphate content.

The distribution of phosphate in the Pacific Ocean (fig. 46) has many features that differ from those in the Atlantic Ocean. As might be expected, the conditions in the Antarctic are rather similar. However, the maximum amounts in the Pacific are found not in the Southern Hemisphere, as in the Atlantic, but north of the Equator, where the amounts present are about twice those found in the Antarctic (3.5 μg-atoms/L and about 2.0 μg-atoms/L, respectively). Furthermore, there is no clearly defined layer of minimum phosphate content beneath the maximum. The deeper waters of the Pacific are in general higher in phosphate than those of the Atlantic Ocean. The difference in the character of the distribution in the two oceans is related to the nature

of the deep-water circulation (p. 754), which is also reflected in the lower dissolved oxygen content of the waters of the Pacific. The latter feature is particularly marked in the Northern Hemisphere.

The amounts of phosphate in the Indian Ocean (fig. 47) are greater than those in the Atlantic but somewhat less than those in the Pacific Ocean. The intermediate maximum in southern latitudes corresponds to that in the Atlantic, and the maximum in the equatorial region corresponds to that in the North Pacific, as it is related to the low oxygen content of the water. The minimum layer, at depths of about 3500 m, is clearly defined in all latitudes north of 40°S.

The differences between the phosphate concentrations in the three oceans are brought out in fig. 48, which is based on data collected by the Dana (Thomsen, 1931) on a voyage around the world and by the Atlantis in the western North Atlantic (Rakestraw and Smith, 1937). Observations from six stations (locations shown in fig. 44) in each ocean have been averaged and the values corrected for salt error (p. 182), so that they are somewhat higher than the values given in the sections. Too great emphasis should not be placed on the absolute values, as the purpose of the illustration is only to show the character of the vertical distribution in the three oceans, and to emphasize the higher phosphate content of the Pacific and the Indian Oceans in contrast to that of the Atlantic Ocean. Data from individual stations in moderate and low latitudes usually show a well-defined intermediate maximum.


Vertical distribution of phosphate in the Atlantic, Pacific, and Indian Oceans based on data from the stations shown in fig. 44.

The lack of nitrate observations from many parts of the sea makes it impossible to prepare sections comparable to those for phosphate. Considerable data have been collected by the Dana (Thomsen, 1931, 1937), the Discovery (Discovery Reports, 1932; Deacon, 1933), the

Meteor in the region to the south of Greenland (Defant et al, 1936) and by the Atlantis, but the information is insufficient to prepare longitudinal sections. In order to demonstrate some of the features of the nitrate distribution, fig. 49 (from Deacon, 1933) has been prepared for the southern portion of the Atlantic Ocean. This corresponds in part to fig. 45 showing the phosphate distribution. As might be expected, there is considerable similarity between the patterns of phosphate and nitrate distribution, although the intermediate maximum is not clearly shown by the nitrate. The Antarctic is extremely high in nitrate. The data from the North Atlantic indicate that the character of the nitrate distribution is very similar to that of the phosphate—namely, that the concentrations are only about one half those present in comparable latitudes in the Southern Hemisphere.


Distribution of nitrate in a longitudinal section in the central Atlantic Ocean. Units: μg-atoms of nitrogen per 20° liter.


Vertical distribution of nitrate in the Atlantic, Pacific, and Indian Oceans based on data from the stations shown in fig. 44.

Curves for the nitrate distribution in the three oceans based on observations of the Dana and Atlantis and comparable to those shown for phosphate are given in fig. 50. Too much emphasis should not be placed upon the absolute values, but the curves clearly demonstrate the characteristic features of the vertical distribution

and the high nitrate content of the Pacific and Indian Oceans in comparison with that in the Atlantic. Data from individual stations in moderate and low latitudes generally show an intermediate maximum somewhere between 500 and 1500 m.

It has repeatedly been emphasized that there is a close parallelism between the concentrations of nitrate and phosphate. This relationship has been demonstrated by plotting against each other in fig. 51 the average data for phosphate and nitrate presented in figs. 48 and 50. It is immediately seen that there is a good linear relationship between the two substances. The straight line represents the “normal” ratio of nitrogen to phosphorus of 15:1 atoms proposed by Cooper (1938a). Because of this relationship, it is possible to predict with a fair degree of accuracy the concentration of either nitrate or phosphate when either one is known, and, as pointed out previously (p. 237), a relationship exists between the concentrations of these elements and the oxygen depletion.


Relation between phosphate and nitrate in the three oceans. Points represent averages for individual depths used in constructing figs. 48 and 50. Straight line represents “normal” ratio proposed by Cooper.

The ranges in the various inorganic forms of nitrogen have been given (p. 181) as


The nitrate is the most abundant form of inorganic nitrogen, and, as shown in fig. 51, the low values occur at and near the surface, the high values in deeper water. The distribution of nitrite and ammonia, which are always in low concentrations, differs from that of nitrate in that the higher values occur in or above the thermocline. Nitrite may also be found near the bottom in shallow water, but it is generally absent from most of the water column. The ammonia content of the deeper waters is relatively uniform and low. Data given by Rakestraw (1936) and by Redfield and Keys (1938) indicate that in deep water, away from shore, the amounts of these substances are small. At one station, half way between Cape Cod and Bermuda, the nitrite was found only at about

75 m, and the ammonia varied irregularly from 0.3 to 0.6 μg-atoms/L between the surface and depths greater than 4000 m. Robinson and Wirth (1934b) found that, off the West coast of Canada and Washington, ammonia was highest on the average near the surface (1.5 μg-atoms/L). Moberg and Fleming (1934) also found the ammonia to be rather irregular off southern California, but independent of depth and present in quantities of about 2.0 μg-atoms/L.


Distribution of silicate in a longitudinal section in the southeastern Atlantic Ocean. Units: μg-atoms of silicon per 20° liter.

The data available on the distribution of silicate in the oceans are even fewer than those for nitrate. The Discovery (Clowes, 1938) has made numerous observations in southern latitudes, and certain of these data have been used in preparing longitudinal sections in the South Atlantic (fig. 52) and in the Indian Ocean (fig. 54). The locations of these sections are shown in fig. 44. The Carnegie obtained numerous observations in the northeastern and central Pacific and there are scattered observations from other regions, but they are inadequate for the preparation of longitudinal sections. In order to bring out similarities and differences in distribution, the silicate sections should be compared with the corresponding phosphate sections (figs. 47 and 53). It is readily seen that the vertical distribution of silicate differs from that of phosphate and nitrate, as there is no marked intermediate maximum and the concentration increases all the way down to the bottom. The reasons for this difference in the pattern of distribution have already been set forth (p. 237). In the Atlantic Ocean (fig. 52) the silicate content of the deeper water is much less in low latitudes than it is in the

far south, but in the Indian Ocean (fig. 54) the contrast is not so great. Data in fig. 55 indicate that the waters of the North Pacific are extremely rich in silicate and contain amounts comparable to those in the Antarctic. The Antarctic region is high in silicate, and also in phosphate and nitrate. Details of the distribution of silicate in the upper layers in the equatorial part of the Pacific Ocean are shown in fig. 198 (p. 710). The Dana made no silicate determinations, and it is therefore impossible to present vertical distribution curves for the three oceans comparable to those for phosphate and nitrate. The high content of silicate in the North Pacific is shown by two curves (Thompson, Thomas, and Barnes, 1934, Barnes and Thompson, 1938) in fig. 55. The amount present below 1000 m (about 170 μg-atoms/L) is somewhat greater than that found in the Antarctic (Clowes, 1938). In order to demonstrate the lower quantities present in the Atlantic and Indian Oceans, curves were constructed from data in fig. 52 at 36°S and in fig. 54 at 2.5°N.


Distribution of phosphate in a longitudinal section in the southeastern Atlantic Ocean. Units: μg-atoms per 20° liter.


Distribution of silicate in a longitudinal section in the western Indian Ocean. Units: μg-atoms of silicon per 20° liter.


Vertical distribution of silicate at individual localities in the North Pacific, South Atlantic, and Indian Oceans.

In basins, the distributions of the elements discussed above may be quite different from those characteristic of the open sea. As shown in chapter IV, the conditions in basins depend upon the topography, the character of the renewal processes below sill depth, and the dissolved oxygen content (aeration) of the water. In well-aerated basins, where there is inflow at the surface, the nutrient content is usually low. For example, in the Mediterranean Sea the phosphate and nitrate below sill depth are small when compared with the concentrations in the waters of the Atlantic Ocean (Thomsen, 1931). In the western Mediterranean below about 1000 m the contents of phosphate and nitrate are constant in amounts of 0.6 μg-atoms/L and 11 μg-atoms/L, respectively, which are about one half or less of the amounts in the open Atlantic. The dense water that flows out of the Mediterranean over the sill and mixes with the intermediate waters of the North Atlantic is therefore relatively low in nutrient elements and tends to reduce the phosphate and nitrate contents of the waters at intermediate levels in the eastern North Atlantic.


In basins of low oxygen content, such as the Red Sea, the nitrate and phosphate are relatively high for reasons already stated (p. 237). In stagnant Norwegian fjords, where hydrogen sulphide is present in the water, Ström (1936) found the phosphate to be as high as 10 μg-atoms/L.

Factors Influencing the Distribution of Nutrient Elements

The concept of dynamic equilibrium (p. 160) can be applied to the large-scale distributions of phosphate, nitrate, and silicate discussed above. On this assumption the distribution is stationary, the local change is zero, (∂s/∂t = 0), and there must be a balance between the effects of diffusion, advection, and the net effects of biological processes. Although the concept of stationary distribution is applicable to the deeper waters and probably to the upper layers in low latitudes, it is not valid for the upper several hundred meters in localities where seasonal variations are conspicuous.

In the surface layers—that is, in the euphotic zone—biological processes will generally lead to a net utilization of the nutrient elements, and, if the rate of utilization exceeds the rate of supply by diffusion and by advection, the concentrations will decrease. This is the characteristic change that occurs during the spring and summer months in regions where physical conditions limit plant activity during the winter. In such regions the reverse change takes place during the winter, when the supply to the surface layers by diffusion and advection outweighs the utilization and leads to an increase in the nutrient content at and near the surface. Consequently in regions where the temperature, light intensity, and biological or other agencies are unfavorable for plant growth during part of the year, marked seasonal variations may occur in the nutrient distribution in the surface layers. In addition to fluctuations in consumption, the supply of nutrient elements to the surface layers by diffusion and advection may vary during the course of the year. For example, the magnitude of the vertical coefficient of eddy diffusivity, Az, is independent of the gradient in the nutrient concentrations but will be affected by the temperature distribution and the wind conditions. Furthermore, changes due to shifts in currents will affect the distributions, although undoubtedly the most conspicuous effects are those related to vertical advection. In regions of surface convergence, waters low in nutrients may extend for a considerable distance below the euphotic layer. On the other hand, where there is divergence—that is, upwelling—nutrientrich waters are carried upward toward the surface. Divergence may occur in the open ocean, as on the Equator and at the northern boundary of the Equatorial Countercurrent (p. 635), or along continental coasts where the prevailing winds are such that upwelling is induced (p. 501). In coastal areas where upwelling is seasonal or intermittent, the nutrient

content of the surface waters may show marked fluctuations and may actually increase during the season of maximum plant activity. In the discussion of the annual cycle of temperature in the upper layers, data were presented for Monterey Bay, California, where temperature changes due to the effects of heat conduction by eddy processes, upwelling, and current shifts could be traced (fig. 32 and p. 131). Corresponding data (Phelps, 1937) or the silicate distribution in Monterey Bay are given in fig. 56. In order to demonstrate the similarity to the temperature changes, the silicate scale in the diagram increases downward. During the spring and summer months, upwelling maintains a virtually constant amount of silicate in the surface layers despite the utilization that must take place, and at a depth of 20 m the silicate actually increases during this period. This seasonal variation is very different from the changes found in the English Channel (see fig. 66) and at Friday Harbor, Washington (see fig. 65), where there are much smaller concentrations of silicate during the summer than during the winter. The shift of the currents at Monterey Bay in September, which brings offshore water toward the coast, is reflected in a sharp drop in the silicate content and a corresponding rise in the temperature of the upper layers. A similar effect is shown in December, when the northward-flowing coastal current is established.


Seasonal variations in the silicate content at various depths in Monterey Bay, California. Silicate scale increases downwards.

From the foregoing comments, it is obvious that seasonal variations in the distribution of the biologically affected elements must be interpreted with great care. Fluctuations in concentration cannot be ascribed to biological processes alone unless it can be established on the basis of other data, such as temperature and salinity observations, that the effects of advection and diffusion can be neglected. In certain regions where there are marked differences between the summer and winter concentrations, estimates of organic production have been made from the depletion of nutrients occurring in the upper layers. Such estimates are minimal unless the effects of regeneration and diffusion are taken into account, and will be invalid if advection is an important factor, as in Monterey Bay. No systematic study of the seasonal variations in the distribution of the nutrients in the open ocean has yet been made, but in certain coastal areas a considerable amount of data has been accumulated that will be taken up under the discussion of the individual elements in the following pages.


Compounds of Carbon, Nitrogen, Phosphorus, and Silicon in the Sea

Organic Carbon. The carbon present as carbon dioxide and the salts of carbonic acid, as well as many of the effects of biological activity on the distribution of CO2, have been discussed in chapter VI. Seasonal changes in the CO2 content of the waters of the English Channel have been described by Cooper (1933). The carbon present in sea water in organic combination will now be considered.

The CO2 that is removed from the water by organisms is utilized partly for the secretion of calcareous structures but chiefly for building up organic compounds. Metabolic activity returns most of the organic carbon to solution as CO2, and bacteria and other microorganisms play an important part in the break-down of excretory products and the detrital material resulting from the death and partial disintegration of plants and animals. If dissolved oxygen is present in the water, the end products will be completely oxidized, but in the absence of oxygen anaerobic bacteria may flourish and hydrogen sulphide and other products of putrefactive decomposition may be formed. The latter conditions apparently occur only in and above the sediments in certain areas and in enclosed basins. The CO2 in calcareous structures returns to solution if the skeletal material dissolves. It should be remembered that, as with other elements, the cycle of carbon in the sea is not completely closed, since there is some loss to the sediments in both calcareous material and in resistant organic matter.

Before proceeding to a discussion of the amounts of carbon occurring in organic combination, either in living organisms or in particulate or dissolved compounds of organic origin, it should be pointed out that the division of the organic matter into various fractions is an empirical one. It has been customary to speak of “net plankton”—usually that which can be removed from the water by filtration through a fine net; “nannoplankton”—that which will pass through the ordinary net but which can be removed by centrifuging or passage through filter paper; and “dissolved organic matter”—that which will pass through the filter. Examination of the literature reveals that a variety of methods have been used to separate those fractions, and consequently the results for different fractions are not always comparable. In the following discussion the term “particulate material” will be used to designate all the material, either living or dead, which is caught by a fine filter that will retain particles of about the size of the larger bacteria. It should be kept in mind that the organisms in sea water, the number of bacteria, and even the inorganic constituents involved in bacterial development undergo rather rapid changes after the collection of samples. Therefore, unless the separation into the required fractions is made immediately, or unless suitable preservatives are added, the results obtained for the

different fractions may be in error. Such errors may account for the relatively low amounts of particulate organic matter sometimes reported.

Pütter (p. 912) maintained that marine invertebrates obtained nourishment from the dissolved organic matter in the water. This hypothesis was based on observations which indicated that the amount of organic matter in solution was many hundred times greater than that present as plankton and particulate detritus. Further investigation has tended to reduce the difference between the two fractions of organic matter, because the earlier determinations of dissolved organic material were obtained by inaccurate methods. Additional studies summarized by Krogh (1931) and Bond (1933) also indicate that dissolved organic matter cannot be utilized by animals. Although Pütter's hypothesis has lost its original significance, it has stimulated a great deal of interest in the problem of the dissolved organic material and its utilization, and investigations have shown that dissolved organic material, although unused by animals, can be utilized by bacteria (p. 912).

Although the problem has attracted much discussion and speculation, there are very few trustworthy data concerning the amounts of carbon present in particulate or dissolved material in the sea. It is extremely difficult to determine accurately the carbon in small amounts of organic matter, especially in the presence of large quantities of salts. Methods have been proposed for concentrating the particulate material by filtration or precipitation (von Brand, 1935) and for determining the carbon by a microcombustion method. No method is as yet available for concentrating the dissolved organic material and freeing it from the salt; consequently the existing methods are based on wet combustions with strong oxidizing agents such as permanganate or chromate. (Bond, 1933; Krogh and Keys, 1934.) Two difficulties are inherent in the latter type of determination: (1) many of the inorganic salts present in sea water interfere with the oxidation and usually tend to give high values, (2) there is uncertainty as to the completeness of the destruction of the organic compounds. Some organic materials may be completely converted into carbon dioxide, water, and so on, by such a procedure; other compounds are only partially decomposed, and still others are not attacked at all. As the chemical constitution of the dissolved material is not known, it is difficult to evaluate the accuracy of determinations made by such methods. Determinations made by wet combustion give the “oxygen consumed,” and a further uncertainty arises when it is necessary to convert these values into the amounts of organic carbon present in the samples.

The development of marine bacteriology has offered a new approach to the problem of the amount of organic material (both particulate and dissolved) in sea water. If sea water is placed in clean, stoppered bottles and kept in the dark, bacteria will develop in great numbers

and the dissolved oxygen will be consumed in metabolic processes. The amount of oxygen consumed, if it is not completely exhausted, is a measure of the amount of organic matter attacked by the bacteria. Even if the water is passed through an ultrafilter to remove particles of colloidal dimensions, and is inoculated with unfiltered sea water, it is found that enough organic matter is still present to permit the development of a large bacterial population. This problem has been discussed by Keys, Christensen, and Krogh (1935) and in a number of publications by Waksman and by ZoBell (for example, Waksman and Renn, 1936; ZoBell, 1940). Although a line of investigation of great promise, studies of bacterial oxygen consumption are beset with many difficulties and the results thus far available are not conclusive. It has been shown that the amount of oxygen consumed is a function of temperature, time, the source of water, and the solid surface-volume ratio (ZoBell and Anderson, 1936). Until standardized methods are established and extensive studies made of the regional, depth, and time variations of this property, only some general quantitative results given below can be considered.

The amount of carbon present in oceanic sea water in inorganic compounds is between 2.1 and 2.5 mg-atoms per liter (25 to 30 mg/L), depending upon the salinity, temperature, and effects of biological activity. Krogh (1931, 1934a,b) has summarized the available data on the amount of organic carbon in sea water. In his later work he reports total organic carbon analyses on six water samples from the Atlantic Ocean. Virtually no variation with depth was found, and Krogh considers the average value applicable to all depths and oceans. The average was 0.2 mg-atoms (2.05 mg) of carbon per liter, which is approximately one tenth of the amount present in inorganic form. From estimates of the amount of plankton, Krogh found the dissolved material to be about three hundred times more abundant than the particulate organic matter. These figures apply to the deeper water of the open ocean. Bond (1933) examined the surface layers in nearshore areas of higher production and found rather different values. His original data, obtained by wet combustion, are expressed in terms of oxygen consumed. In order to make them comparable to those of Krogh, it has been assumed that two atoms of oxygen were required to oxidize one atom of carbon. The recomputed minimum, maximum, and average values are given in table 53. Although Bond's values for the dissolved fraction are approximately the same as those of Krogh for the total carbon, they show a considerable range. Furthermore, it will be noted that the particulate material is relatively more abundant and forms between one tenth and one third of the total.

The results of bacteria1 oxygen-consumption studies are difficult to evaluate because of the variety of techniques that have been used.

Sometimes the water has been filtered and at other times it has not, and in relatively few cases have the cultures been kept for a sufficiently long time. The maximum values of oxygen consumption for unfiltered sea water range between 0.13 and 0.18 mg-atoms of oxygen per liter (1.5 and 2.0 ml/L), Using a 2 : 1 ratio by atoms, these are equivalent to 0.07 to 0.09 mg-atoms of carbon. These values are of the order of one fifth to one half of the total organic carbon values given above. Waksman and Renn (1936) have found that in the laboratory about 50 per cent of the organic matter is readily attacked by bacteria, of which about 60 per cent is oxidized and 40 per cent is converted into bacterial cell substance. Estimates of organic carbon in sea water obtained in this way therefore give values of the same magnitude as those obtained by chemical methods.

Substance mg-atoms/L of carbon
Minimum Maximum Average
Net plankton 0.008 0.06 0.03
Nannoplankton 0.005 0.11 0.03
Dissolved 0.13 0.25 0.20
Total organic 0.143 0.42 0.26

The oxygen consumptions given above were for water from near the surface in areas relatively rich in plankton. Samples from deeper levels consume about one half as much oxygen. As shown previously (p. 236), there are relatively constant ratios between carbon, nitrogen, and phosphorus in the organic material. This fact has been used in certain studies in which the organic nitrogen (determined by the Kjeldahl method) is used as a measure of the amount of organic matter. The relative amounts of organic nitrogen and phosphorus in sea water are in fair agreement with the amounts of carbon given above. The organic nitrogen content of bottom samples (p. 1010) has been widely used as a measure of their content of organic matter. In sediments the ratio of

carbon to nitrogen has been found to be larger than in the organisms. This change in the ratio indicates that a relatively large proportion of the nitrogen has been lost by the refractory detrital matter that accumulates on the sea bottom.


Annual cycle in the nitrate and nitrite content of the surface waters at Friday Harbor, Washington.


Seasonal variations in nitrate, nitrite, and ammonia in the surface layer (0–25 m) and in the bottom layer (50–70 m) in the English Channel during the period November, 1930, to January, 1932. (After Cooper, 1937b.)

Nitrogen Compounds and Their Seasonal Variation. In certain coastal areas, sufficient data are available to examine the seasonal changes in the distribution of nitrate, nitrite, and ammonia. Only selected cases of seasonal variations will be given, but additional references may be found in the works cited. Phifer and Thompson (1937) give the results of nearly five years' studies of the surface conditions at Friday Harbor, on the San Juan Channel. The averages of the monthly means for NO3 and NO2 for the period 1931 to 1935 are shown in fig. 57. It should be noted

that the NO2 scale is one fiftieth that for the NO3. The nitrite appears in greatest abundance after the period of most rapid utilization of nitrate by plants in the spring and summer, and then decreases and is minimal at approximately the period of maximum nitrate. Cooper (1937b) has shown the seasonal changes in NO3, NO2, and NH3 in the surface (0 to 25m) and bottom (50 to 70m) layers in the English Channel during the interval November, 1930, to January, 1932 (fig. 58). The three components are on different scales—namely, NO3-N: NH3-N: NO2-, N = 8 : 2 : 1. In general, these data show cycles similar to those in fig. 57. During and after the period of greatest plankton development, there is a rise in ammonia, followed by one in nitrite and then one in nitrate. This indicates that, in the regeneration of nitrate from organic matter, the nitrogen passes through these stages. It should be noted, however, that the ammonia and nitrite never reach concentrations as great as the nitrate. In the English Channel the total inorganic nitrogen compounds are always much lower than at Friday Harbor. Rakestraw (1936) has presented detailed observations of the variations in nitrite and nitrate during a year in the Gulf of Maine, from which fig. 59 is taken. These data show the nitrite to be most abundant near the surface during the summer and autumn, when the nitrate is lowest. That the higher quantities of nitrite are definitely associated with the distribution of density, and hence of temperature, is shown in fig. 60 (Rakestraw, 1936). When there is a marked thermocline, the nitrite is either in or above it. Similar data are not available for ammonia in the Gulf of Maine, but Redfield and Keys (1938) report that it is closely related to the amount of nitrite and also to the amount of plankton in the water.


Seasonal variations in the vertical distribution of nitrite and nitrate in the Gulf of Maine during the period May, 1933, to September, 1934. (After Rakestraw, 1936b.)

The nitrogen in particulate organic material may be determined on the separated material which has been concentrated by filtration or

carried down with a flocculent precipitate (von Brand, 1935). von Brand (1938) has determined the particulate organic nitrogen for five oceanic stations in the northwest Atlantic. The greatest variability was found in the upper 400 m, with values ranging between about 0.07 and 1.3 μg-atoms/L. The high values usually occurred at or near the surface. Near Iceland, values as high as 5.2 μg-atoms/L of nitrogen have been found, and in the Gulf of Maine surface values of 2.4 μg-atoms/L were obtained (von Brand, 1937). Below 400 m the amounts varied rather irregularly between 0.07 and 0.21 μg-atoms/L. Cooper (1934) found between 0.3 and 0.7 μg-atoms N/L as net plankton in the English Channel. Cooper's samples did not include nannoplankton and detritus.


Vertical distribution of nitrite, as related to density (σt) and temperature. (After Rakestraw, 1936.)

The total organic nitrogen, including both particulate and dissolved material, has been investigated by Robinson and Wirth (1934a,b). Kjeldahl analyses on unfiltered oceanic sea water showed about 7.2 μg-atoms/L of organic nitrogen near the surface, about half this amount at intermediate depths, and a slight increase again toward the bottom. In nearshore water the values near the surface were about twice as high as those in the oceanic samples. Moberg and Fleming (1934), using a similar method, found about 10 μg-atoms/L, on an average, of organic nitrogen in the surface layers off southern California and somewhat higher values at greater depths.

Figures 57 and 58 show that at Friday Harbor about 10 μg-atoms/L of NO3-N disappear during the summer, and that the change in the English Channel is approximately the same. As NH3 and NO2 are rarely present in comparable amounts, we must conclude that the nitrogen is in organisms, organic debris, dissolved organic compounds, or in some unrecognized inorganic form.

Interesting experiments by von Brand, Rakestraw, and Renn (1937, 1939) on the regeneration of nitrate from marine plankton in vitro indicate that the formation of ammonia from organic matter probably takes place without the formation of intermediate compounds. The results

of one of their experiments are shown in fig. 61. In this experiment, sea water, with added diatom material, was placed in the dark. After about four months, much of the particulate-N had been converted to NO3 through the intermediate stages of NH3 and NO2. The jars were then placed in the light and inoculated with diatoms. Almost complete utilization of the NO3 followed. The jars were again placed in the dark and the cycle was repeated.

These experiments are extremely interesting, although regenerative processes in vitro apparently differ considerably from those in the sea. This result might well be expected from the peculiar laboratory conditions and the fact that the water was enriched with organic material. It may also account for the fact that the NH3 and NO2 reached relatively high values (the same as the NO3), and that during the regeneration the stages of production of NH3, NO2, and NO3 were clearly defined (cf. figs. 57 and 58). Furthermore, it is interesting to note the great increase in the amount of nitrogen present as particulate material that was apparently resistant to the action of the bacteria present. At the end of the experiment, approximately 50 per cent of the nitrogen was in this form. Such a “waste” of nitrogen does not take place in the sea. It is also of interest to note that diatoms would flourish if the medium were placed in the light when either NH3 or NO2 were abundant and before the NO3 had been produced. This supports the theory that marine plants can use any of these inorganic forms bf nitrogen equally well.


Experiment on the utilization and regeneration of nitrate. When medium was placed in light, it was innoculated with diatoms. Data from von Brand, Rakestraw, and Renn, 1939.

Nitrogen Cycle in the Sea. The chemically bound nitrogen in sea water is known to occur in living organisms, in dissolved and particulate material of organic origin, and as ammonia, nitrite, and nitrate.

The quantities present in these various forms vary from place to place, and in the upper levels may undergo seasonal changes. Much work has been done on the nitrogen cycle in the sea to determine the forms of nitrogen that can be used by the plants and the agencies which return the organic nitrogen to inorganic forms. In the surface layers, prior too the vegetative season, the most abundant inorganic form of nitrogen is nitrate, and in deeper water, where nitrite and ammonia are negligible, this is always the case.

Bacteria play important roles in the mineralization of organic nitrogen by acting upon detrital material, excreta, and dissolved organic matter, and in the intertransformation of NH3, NO2, and NO3 (chapter XVIII). Pure culture studies with different species of marine bacteria show that a variety of transformations can be made under laboratory conditions, but these observations must be applied to the sea with caution because the organisms may not be capable of carrying on similar processes in the natural environment. The very large literature bearing on the nitrogen cycle in the sea has been reviewed by Cooper (1937b), who has considered all the various modes of transformation that are possible and from them has selected the more probable ones. In general, it is believed that the nitrogenous material gives rise to ammonia, which in turn is converted to nitrite and then to nitrate. The ammonia may be formed by the hydrolysis of protein material, amino acids, amines, and purine compounds such as urea, or through bacterial action on them.

The oxidation of ammonia to nitrite releases a large amount of energy and hence needs only to be activated in some way. The following agencies have been suggested:

  1. Photochemical oxidation induced by direct sunlight. This reaction was first observed in sea water by ZoBell (1933), but, as pointed out by Cooper, can be effective only within the upper meter or so of water, owing to the rapid absorption of the shorter wave lengths that activate the reaction.

  2. Chemical oxidation by the free oxygen in the water in the presence of surface catalysts. This reaction is of unknown significance.

  3. Bacterial oxidation. Nitrifying bacteria are present in bottom sediments, and forms isolated by Zobell (1935b) converted ammonia to nitrite. However, the conversion occurred at a much higher oxidation-reduction potential than is ordinarily found in the sediments. Studies of the decomposition of marine plankton in vitro show a conversion of ammonia to nitrite, but no nitrifying bacteria could be isolated (von Brand, Rakestraw, and Renn, 1937). Nonetheless, the fact that nitrifying bacteria cannot be readily detected in sea water is not definite proof of their absence. It is well known that many marine bacteria are difficult to culture, and development of new techniques may establish their presence. Carey (1938) has shown that they can be isolated from

    water rich in plankton. The presence of nitrite, which can sometimes be detected near the sea bottom, as in the English Channel in certain seasons, may indicate either oxidation of ammonia or reduction of nitrate. Although it has not been established, bacteria are probably the most important agency in the oxidation of ammonia to nitrite.

Oxidation of nitrite to nitrate also releases energy, and, as in the oxidation of ammonia, purely chemical or photochemical processes may be important. Bacteria capable of making the transformation are abundant in sediments, but they are difficult to isolate from the water column. Development of suitable techniques may also clarify this problem. Cooper has pointed out that in sea water saturated with oxygen the nitrate in equilibrium with nitrite will be of a tremendously greater order of magnitude. Hence, the detection of nitrite in the water column may indicate active production of this substance, which is present only as a transitional stage in the regeneration of nitrate. Brandt's hypothesis (p. 768) was based on the discovery of marine bacteria which under laboratory conditions were capable of denitrification. However, it is now considered that under the conditions prevailing in the sea there is little or no loss of fixed nitrogen, although it has been shown that reduction of NO3 to NO2 may occur. This reduction may also be carried out by diatoms, as has been observed in pure-culture experiments (ZoBell, 1935a).


Annual cycle in the phosphate content of the surface waters at Friday Harbor, Washington, and the monthly N/P ratios.

Organic Phosphorus and Seasonal Variations in Phosphate. For certain coastal areas, sufficient data exist to show the nature of the seasonal variations in phosphate. Friday Harbor, the English Channel, and the Gulf of Maine have been selected as examples.

In fig. 62 are presented the averages based on about four years' observations of the monthly mean values for PO4-P at Friday Harbor (Phifer and Thompson, 1937). Highest values occur during the winter, and lowest values during the summer season, when phytoplankton growth has been great. The monthly ratios of N:P are somewhat lower than the normal ratio proposed by Cooper, and their variability during the course of the year indicates that proportionally more nitrate than phosphate is utilized. Neither of these substances can be considered as limiting the amount of phytoplankton produced at this locality. Phosphate data from the English Channel (Cooper, 1938b) for the period November, 1930, to January, 1932, are presented in fig. 63 as

PO4-P. In this shallow region there is never very much difference between the quantities in the surface and in the bottom layers. The amount of phosphate in the English Channel is much less than at Friday Harbor, although the difference between maxima and minima are of the same order, 0.5 μg-atoms/L in the former, and about 0.75 μg-atoms/L in the latter area. The minimum values for San Juan Channel are in excess of the maximum values for the English Channel. Similar conditions hold for the nitrate (p. 252)


Seasonal variations in the phosphate content of the surface layer (0–25 m) and the bottom layer (50–75 m) in the English Channel during the period November, 1930, to January, 1932.

Cooper (1938b) has assembled phosphate data for the English Channel covering a period of eighteen years. In fig. 64 are entered the winter maxima for the average PO4-P content of the water column near Plymouth. As the phosphate content of the water during the winter is a measure of potential production for the following spring and summer, his data indicate a drop in fertility after 1929 or 1930. The changes are considered to be associated with the circluation, which may undergo random or periodic fluctuations.


Winter maxima for phosphate in the English Channel for the period 1921–1938.

In the preceding discussion, emphasis has been placed on the cyclic nature of the seasonal variations, and it is obvious that the conditions in any area may not repeat themselves if disturbances, such as shifts in the circulation, bring about changes.

The seasonal variations in the PO4-P and in the various organic phosphorus fractions in the Gulf of Maine have been studied by Redfield, Smith, and Ketchum (1937). During one year, five series of samples were collected at various depths between the surface and the bottom. These

samples were analyzed for PO4-P, dissolved organic-P, and particulate organic-P. The results are summarized in table 54, The variations in total phosphorus during the year are ascribed to the fact that, although collections were made at the same locality, the circulation (advection) brought in water of different character. Table 54 shows that the particulate phosphorus never represents more than about 10 per cent of the total, but that the dissolved organic phosphorus sometimes occurs in relatively large amounts, in the upper 60 m approaching 50 per cent of the PO4-P.

Form of phosphorus Interval of depth, meters μg-atoms/L of phosphorus
May 18, 1935 Aug. 20, 1935 Nov. 8, 1935 Feb. 26, 1936 May 14, 1935
Phosphate 0–60 0.60 0.68 0.65 1.03 0.64
60–120 1.11 0.91 1.08 1.02 1.25
120–180 1.31 1.22 1.25 1.11 1.51
180–240 1.61 1.39 1.22 1.51 1.60
Dissolved organic 0–60 0.08 0.34 0.29 0.07 0.14
60–120 0.02 0.29 0.31 0.14 0.36
120–180 0.01 0.17 0.29 0.17 0.15
180–240 0.00 0.20 0.37 0.03 0.10
Particulate organic 0–60 0.15 0.10 0.10 0.05 0.12
60–120 0.06 0.05 0.05 0.05 0.07
120–180 0.04 0.03 0.08 0.03 0.02
180–240 0.04 0.03 0.08 0.04 0.06
Total (Average for whole water column) 0–240 1.26 1.36 1.44 1.31 1.51

Earlier work on organic phosphorus in the sea has been summarized by Cooper (1937a), who also reports observations from the English Channel. Cooper points out that many determinations of the total “organic phosphorus” (both particulate and dissolved) probably include arsenite-arsenic that has been oxidized to arsenate when the organic matter was destroyed and that has not been reduced again. Arsenate will give the same reaction as phosphate in the colorimetric estimation, and, because the arsenic present as arsenite will not affect the inorganic phosphate analyses, the determination of “organic phosphorus” will be too high by about 0.2 μg-atoms/L (the concentration of arsenic). In the determinations in the Gulf of Maine this source of error was

eliminated. The “organic phosphorus,” uncorrected for arsenic, averages about 0.4 μg-atoms/L (Cooper, 1937a). If reduced by one half to eliminate the effect of arsenic, this value corresponds to the average for the Gulf of Maine.

The maximum values reported by Redfield et al in the Gulf of Maine for particulate and dissolved organic phosphorus are

If multiplied by 16, the ratio of N:P in organisms, the corresponding values of particulate and dissolved organic nitrogen would be 3.4 and 9.3 μg-atoms/L respectively, which agree with determinations of nitrogen in the corresponding fractions (p. 254).

Phosphorus Cycle in the Sea. The cycle of phosphorus in the sea is rather similar to that of nitrogen except that only one inorganic form, phosphate, is known to occur. As shown above, phosphorus can be found in organisms, in particulate and dissolved organic compounds, and as phosphate. Probably only the phosphate is utilized by plants, and the dissolved organic fraction, which can originate as a metabolic product and from excreta, and the decomposition of organic material must be intermediate stages in the regeneration of phosphate. The roles that bacteria play are not yet known.

Studies of the decomposition of plankton material in vitro have led to some interesting results. Cooper (1935) added zooplankton and phytoplankton material of known phosphorus content to sea water and determined the rate at which the PO4-P was formed. The PO4 appeared more rapidly in the zooplankton samples than in the diatom material. Furthermore, the PO4 produced in the zooplankton samples was in excess of that originally present as PO4 plus that added in the particulate material. This excess was formed from dissolved organic phosphorus originally present in the water. The PO4 in the diatom experiments never rose to the level of the original PO4 plus that added, even after an interval of about five months. The initial rate of appearance of PO4 was rapid, and in the zooplankton experiments the transformation was nearly complete in about two weeks. The difference in behavior of the plant and animal material cannot yet be accounted for. Seiwell and Seiwell (1938) found that zooplankton decomposition at 22° to 25°C (probably higher than the temperature of Cooper's experiments) was such that the formation of PO4 was most rapid during the first day or two after death. This stage was often followed by a period when the rate of utilization of PO4 by microorganisms in the experiments exceeded that of formation. On the basis of these experiments it is sometimes considered that the PO4 must be regenerated more rapidly than the NO3. However, field evidence offers little support for this theory. Examination

of figures illustrating the seasonal variation shows that the minima and maxima in NO3 and PO4 occur in the same months, which would indicate that the relative rates of utilization and regeneration are about the same. Furthermore, estimates of plankton production in the English Channel based on the drop in PO4 and NO3 from winter maxima to summer minima (Cooper, 1938b) yield approximately the same results.
If these are utilized in the proportion of 16:1, the NO3-N equivalent to the PO4-P would be 7.4 μg-atoms/L. As only a small fraction of these amounts are ever present in living organisms, the water must contain an abundance of detrital or dissolved substances available for regeneration. If the rate of regeneration of PO4 were quicker than that for NO3, it would be expected that the drop in PO4 would be much smaller. Although nothing is yet known concerning the absolute rates of regenera- tion in the sea, at least there is no evidence that their relative rates are very different.

Redfield, Smith, and Ketchum (1937), on the basis of the material summarized in table 54, have calculated the manner in which the various phosphorus compounds vary in the different layers, assuming (1) that all utilization of phosphate takes place in the upper 60 m, (2) that all downward transport of phosphorus is due to the settling of organisms and particulate material, and (3) that the upward transport is all due to eddy diffusion. From their examination it was concluded that decomposition and regeneration took place throughout the column of about 240 m of water.


Annual cycle in the silicate content of the surface waters at Friday Harbor, Washington.

Seasonal Variation in Silicate. Observations on the seasonal cycle in silicate-silicon have been carried out in many localities, of which Friday Harbor and the English Channel have been selected as examples. The data presented are comparable to those given for NO3-N and PO4-P. Fig. 65 shows the average monthly values for Friday Harbor based on more than four years' observations. The highest values occur during the winter and the lowest values during the early summer. The range in silicon is 15 μg-atoms/L. The ranges in SiO3-Si, NO3-N, and PO4-P are in the proportions of 20:14.7:1. Data for the English Channel (Cooper, 1933) are shown in fig. 66. It will be noted that the concentration of silicon in the English Channel is much lower than it is at Friday Harbor, being in general only about 1/25 as great. The range in silicon

during 1931 was about 3.5 μg-atoms/L, and the ratios of the ranges of SiO3-Si, NO3-N, and PO4-P are 7.6 : 13.7 : 1. The ratios of utilization of NO3 and PO4 in the two areas are in reasonably good agreement, but there is a marked difference in the amount of silicon withdrawn. Whether or not this difference is associated with the character of the plankton or the sequence of plankton development in the two areas is not known. The seasonal cycle in the SiO3 distribution in Monterey Bay was presented in fig. 56.


Seasonal variations in the silicate content of the surface layer (0–25 m) and the bottom layer (50–70 m) in the English Channel during the period November, 1930 to January, 1932.

A factor that may complicate the seasonal changes in the silicon concentration in nearshore areas is the amount of siliceous material carried in by river waters, which is generally several times greater than that found in sea water; dilution will therefore tend to raise the concentration of silicate (Hutchinson, 1928).

The Silicon Cycle in the Sea. The depletion of the silicon in the surface layers is the result of biological activity and the sinking of the organisms or of their skeletal remains. The silicon removed from the water by diatoms or other organisms may return to solution after the death of the organism or may be deposited on the sea bottom. Unlike their roles in the cycles of nitrogen and phosphorus, bacteria are probably not directly involved in the re-solution of silicon from skeletal material. The skeletal material that sinks to the bottom forms either a temporary or permanent constituent of the sediments. Siliceous sediments are found in higher latitudes where the bottom material contains a very large proportion of diatom frustules, while in other regions radiolarian skeletons make up a large part of the sediment (chapter XX).


In the English Channel the silicon distribution is more erratic than that of PO4 and of NO3, and there may be large differences in the amounts at various levels. High values often occur near the surface and immediately over the bottom. The high surface values are ascribed to the effects of river water, the high bottom values to the re-solution of skeletal material that has settled there. The accumulation of dissolved silicates at a marked thermocline is also evidence of re-solution of slowly settling debris.

Estimates of plankton production in the English Channel based on the difference between the winter maxima and summer minima in PO4-P, NO3-N, and SiO3-Si (Cooper 1933, 1938b) yield values from the SiO3 data which are approximately 1/15 of those obtained from the other elements. Cooper has attributed the smaller value to the rapidity with which the silicon passes through its cycle. The difference in the amount of silicon utilized in the English Channel and at Friday Harbor was noted above, and the fact that the concentration is reduced in the English Channel to a very low level may indicate that in that locality it does influence the production of diatoms.

King and Davidson (1933) found that the quantity of SiO3-Si in solution affected the growth of diatoms in laboratory cultures. They also followed the changes in dissolved silicon after the death of the diatoms and found that complete solution took place in about five months. Marine phytoplankton were similarly studied, and it was found that samples which were boiled dissolved more slowly than those unboiled. The authors suggest the possible existence of an enzyme that hastens solution.


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VIII. The Sea as a Biological Environment

In the foregoing chapters an account has been given of the chemical and physical aspects of the elements that together constitute the inorganic marine environment—namely, (1) the sea water itself, and (2) the ocean floors. The chemical constituents and the physical properties of the sea water, together with their distribution, concentrations, and cyclic changes, the movement of the water, and the nature of the ocean floors are decisive factors in the history and fate of a perplexing array of living things. Herein are held many secrets of racial development, and herein must be sought the understanding of the delicately balanced maintenance of life and of the potentialities of future development.

Marine organisms are to be considered a part of the sea as it exists today. Just as sea water includes the various salts, both conservative and nonconservative (biologically changed), so also it includes the multitude of organisms which are bound to the sea for their existence and which, by origin, are a part of the sea both racially and individually. The organisms, like the saIts, are subject to the natural laws of the sea and are a part of the perpetual cycle of inorganic and organic substances so important in many aspects of oceanography. The changes that are apparent in concentration represent only patterns that are inherent in the phases of the cycle or that result from other causes, such as currents and processes of mixing.

The aquatic environment offers the greatest intimacy between itself and the organisms which it bathes both over the body surface and within open or partially closed cavities as, for example, the internal systems of coelenterates, echinoderms, and tunicates. Because of the stability of the physical characteristics of the sea water and of the composition and concentration of the dissolved salts the organisms, in general, have not developed highly specialized integuments and regulatory systems to protect themselves against sudden and intense environmental changes, as have most land animals. It follows that small changes in the aquatic medium are promptly brought to play upon its population. It should be borne in mind, also, that the organisms themselves, being a part of the dynamic environment, modify particularly its chemical character by withdrawing or adding substances associated with the activities of life.


In subsequent chapters we shall discuss the relation of some of the measurable environmental factors to such phenomena as distribution, propagation, survival, and special adaptations, but first some facts of general application must be considered.

Physical and Chemical Characteristics of the Marine Environment

Water is essential to the maintenance of all life. It constitutes 80 per cent or more by weight of active protoplasm. It is the most efficient of all solvents and carries in solution the necessary gases, oxygen and carbon dioxide, as well as the mineral substances necessary to the growth of plants and animals, and it is itself one of the essential raw materials in the manufacture of foods by plants.

Organisms living in the terrestrial environment have devised means, such as impervious integuments, to conserve water, and the land plants have roots and special vascular systems for transport of water to all growing parts. In the marine environment there is freedom from dessication, except at high-tide levels, and therefore no highly specialized means are provided for conservation of water or for its transport in plants.

Also of biological importance are the high heat capacity of water and its high latent heat of evaporation, both of which obviate the danger that might result from rapid change of temperature in the environmental medium. Owing to the high degree of transparency of water it is possible for the sea to sustain plant life throughout a relatively deep layer, and in animals the development of organs of vision and of orientation has progressed to a marked degree.

Sea water is a buffered solution; that is, changes from acid to alkaline condition, or vice versa, are resisted (p. 195). This property is of vital importance to the marine organisms, mainly for two reasons: (1) an abundant supply of carbon can be available in the form of carbon dioxide for the use of plants in the synthesis of carbohydrates without disturbance to the animal life that may be sensitive to small changes in pH, and (2) in the slightly alkaline habitat the many organisms that construct shells of calcium carbonate (or other calcium salts) can carry on this function much more efficiently than in a neutral solution.

The support offered to the bodies of marine organisms by the specific gravity of the surrounding medium obviates the need of special supporting skeletal structure in many forms. Striking examples of these are the jelly fishes, unarmored molluscs, unarmored dinoflagellates, and even the large marine mammals with their heavy skeletons, which could not survive in their present bulky state except in an aquatic habitat. The hard shells of crabs, clams, snails, and so on, doubtless serve as support, especially in some burrowing and intertidal forms, but these hard parts may be looked upon also as protective and as a framework for attachment of muscles used in digging, creeping, or swimming.


Sea Water and the Body Fluids. Sea water is a most appropriate environment for living cells, since it contains all of the chemical elements essential to the growth and maintenance of plant and animal protoplasm. It has been shown that sea water is a solution of a large number of salts, and it is important here to consider how it is related as an external fluid medium to the “internal medium”—namely, the body fluids (blood, coelomic fluid, and so on) of the organisms. The ratios of the major salts to each other, and usually their total concentration also, are strikingly similar in sea water and in the body fluids of marine invertebrates. The similarity of composition is not confined to marine animals, however, but is also in evidence in modified form in both terrestrial and fresh-water animals, including the lower and higher vertebrates, as is shown in table 55, which is from data compiled by Pantin (1931) and expanded by Dakin (1935).

Osmotic Relationships.It is well known that when solutions of different osmotic pressure are separated by a semipermeable membrane that allows the passage of water but not of the solutes, there is a movement of the water through the membrane into the more concentrated solution. The cell membranes of organisms are just such semipermeable membranes through which a movement of fluids occurs inward or outward, depending upon whether the osmotic pressure of the external medium is less (hypotonic) or greater (hypertonic) than the internal medium. The internal and external media are isotonic when they are of equal osmotic pressure.

The osmotic pressure of a solution can be computed from the freezing-point depression (p. 67). This computation is possible because the salts that increase the osmotic pressure of a solution also depress its freezing point. The freezing-point depression below 0°C has been designated by Δϑ, (p. 67), but will here be abbreviated to Δ. Sea water having a salinity of 35.00 ‰ freezes at − 1.91°, owing to depression by the substances in solution. In other words, the value for Δ is 1.91°. Similarly, we obtain a Δ of 0.56 for human blood with a freezing point of −0.56°C.

On the basis of Δ values, the osmotic relations of the body fluids of marine and fresh-water animals to their external environmental medium are compared in table 56, from data compiled by Dakin (1935), to whose review the reader is directed for much greater detail and historical treatment.

From the few examples in the table it is evident that the body fluids of marine invertebrates are isotonic or nearly so with their fluid environment, whereas in the fresh-water forms the body fluids are hypertonic to the dilute external medium. For this reason the marine environment in its osmotic relations fails to exact of its inhabitants as great an expenditure of energy in maintaining the proper concentration of body fluids as does the fresh-water environment. The exact mechanism whereby the fresh-water animals are independent of the external medium and are able to maintain a homoiosmotic condition (that is, steady value for Δ) in the presence of the hypotonic water is not known (see Δ for the eel Anguilla anguilla in fresh and salt water, table 56). Their existence under these conditions, however, requires a constant expenditure of energy in eliminating, through the kidneys and other excretory organs, the excess water taken in by osmosis. Marine invertebrates are poikilosmotic (Δ changing with that of the external medium) only within rather narrow limits (Dakin, 1935); hence, they, too, must have some regulating mechanism. Except in estuarine conditions, however, the range of salinity in most parts of the sea is perhaps within the limits of poikilosmoticity of the invertebrates living there. For example, the lugworm, Arenicola marina, in Helgoland waters with a Δ1.72 has an internal medium Δ1.7, but in the Baltic Sea with a water of Δ0.77 the same species has a Δ value of 0.75 for the internal medium.

Date taken or calculated from Na K Ca Mg Cl SO4
Sea Water 100 3.6 3.8 12.1 180 25.2
Aurelia flavidula (mesogloea) Macallum (1926) 100 5.2 4.1 11.4 186 13.2
Limulus polyphemus Macallum (1926) 100 5.6 4.1 11.2 187 13.4
Aplysia limacina Bethe (1929) 100 4.0 4.4 11.0 180
Homarus americanus Macallum (1926) 100 3.7 4.9 1.7 171 6.7
Acanthias vulgaris Macallum (1926) 100 4.6 2.7 2.5 166
Carcinus maenas Bethe (1929) 100 4.8 4.5 4.8 180
(Cod) Gadus collarus Macallum (1926) 100 9.5 3.93 1.41 149.7
(Pollock) Pollachius virens Macallum (1926) 100 4.33 3.10 1.46 137.8
Frog Macallum (1926) 100 11.8 3.17 0.79 135.6
Dog Macallum (1926) 100 6.6 2.8 0.76 139.5


It should be mentioned here that the teleost (bony) fishes in marine waters are definitely hypotonic and, therefore, in order to keep their body fluids down to the required osmotic pressure for the species, they secrete chloride through the “chloride cells” of the gills (Keys, 1933). This function is a regulation toward a low osmotic pressure of the blood, as opposed to regulation toward a high one as performed by the kidneys of animals in fresh-water environments. That this group of aquatic animals has achieved a marked degree of independence of the osmotic pressure of the external medium is evidenced especially by such forms as the salmon and eel, both of which, though practically homoiosmotic, spend their lives partly in hypotonic and partly in hypertonic environments, The elasmobranchs—namely, the sharks and rays—are isotonic with sea water, but in these the high osmotic pressure of the blood is due not only to the presence of such salts as occur in sea water, but also to high urea content. For further discussion of salinity as an environmental factor, see also p. 839.

Other Characteristics of the Environment

In addition to the chemical and physical properties of sea water, certain other biologically important characteristics are inherent in the marine environment as a whole. These result from the magnitude of the ocean itself, its great depth, and its expanse.

In considering the ocean in its entirety as an environment, we are at first impressed by the wide ranges of living conditions, the salinities varying from those of dilute estuarian waters to concentrations of 37 ‰ or more in the open sea, temperatures from 30°C to freezing point, light intensities from brilliant sunlight at the surface to absolute and perpetual darkness in the deeper layers, and pressures from a single atmosphere at the surface to about 1000 atmospheres in the greatest oceanic deeps.

Marine Animals
Species Internal medium Δ°C External medium Δ°C

* Dakin includes certain observations from the region of Naples giving the Δ of the sea water as 2.29°. These observations, which are often quoted in the literature, have been omitted here because a Δ of 2.29° corresponds to such a high salinity (43.5 ‰) that it must be in error. The maximum salinity in the western Mediter-ranean is about 39 ‰ and the corresponding Δ is 2.14°.

Arenicola marina 1.72 1.7
Arenicola marina 0.77 0.75
Ostrea edulis 2.23 2.11–2.14
Mytilus edulis 2.26 2.11–2.14
Octopus vulgaris 2.16 2.11–2.14
Homarus americanus 1.82 1.80
Cancer pagurus 1.84–1.91 1.91
Hyas aranea 1.83 1.80
Limulus polyphemus 1.90 1.82
Ascidia mentula 2.08 1.98
Teleost fishes
Pleuronectes platessa 0.787 1.9
Conger vulgaris 0.77 2.14
Gadus aeglefinus 0.74 1.92
Fresh-water Animals
Anodonta cygnea 0.09
Unio pectorum 0.15
Limnaea stagnalis 0.22–0.23 0.02–0.03
Telphusa fluviatile 1.17
Daphnia magna 0.20–0.67
Potamobius astacus 0.80
Eriocheir sinenais 1.09
Astacopsis 1.1
Teleost fishes
Salmo fario 0.57
Anguilla anguilla 0.62
(in sea water) 0.73 1.87
Barbus fluviatilis 0.50
Cyprinus carpio 0.50
Anabas tetudineus 0.64
Dipnoi fishes
Epiceratodus fosteri 0.42


Impressive as these ranges may be, nevertheless very uniform conditions do prevail over extensive areas of the environment, and many organisms may, by reason of the monotony of these extensive areas, be very delicately attuned to the prevailing unvarying conditions. Hence, it follows that faunal areas characterized by specific forms can be recognized. On the other hand, a wide range of conditions may be encountered in more restricted areas, especially in coastal regions. These conditions may be due to the physiographic character of the coastline, depth to bottom, topography and nature of the bottom, inflow of land drainage, meteorological conditions, and so forth. Specially adapted and tolerant forms occur here in profusion, for, as will be shown in later chapters, the shallow depths and varying conditions are frequently favorable to abundant production of primary food.

It must not be overlooked that the gradients of salinity, light, and temperature that exist in the sea are favorable to a number of sensitive animals that possess the ability, through swimming or otherwise, to adjust themselves to optimum conditions.

Depth and Light. Inherent in the vertical range or depth of the open-sea habitat are a number of important features of far-reaching biological effect. Of prime importance is the relatively great vertical range of the euphotic zone available for production of floating microscopic plants. But the gradient of light, both as to quantity and quality, resulting from depth of water also allows adjustment of many animals to the optimum condition with respect to this factor and, indeed, is associated with diurnal migrations of many forms to lighter or darker situations.

Pressure. Pressure in itself does not exclude life from the abyssal regions of the sea, for water is but little compressed and equilibrium exists between the inner and outer pressure affecting the body tissues. However, pressure may limit the vertical range of motile forms, although some eurybathic animals apparently are not seriously affected and are known to make daily vertical wanderings of up to 400 m, corresponding to pressure variations up to 40 atmospheres. Harpooned whales are said to “sound” to a depth of 800 m, and the sperm whales must descend normally to great depths, since the large squids upon which they feed inhabit very deep water.

Water Movements. The sea must be viewed as an environment that for the most part is in constant motion with both regular and irregular patterns of flow. The principal biological benefits derived from the circulation are (1) oxygenation of subsurface water, (2) dispersal of wastes resulting from processes of metabolism, (3) dispersal of plant nutrients and other variable elements essential to plant and animal

growth, and (4) dispersal of spores, eggs, larvae, and also many adults. On the whole, the circulation of water is of direct benefit, yet instances may be noted where some adverse conditions result either as incidental or as permanent features. Incidental disturbances may be due to unseasonable shifts in the regular current system, such as give rise to the appearance of “El Niño” off the west coast of South America (p. 704). In this instance, warm water of the Equatorial Countercurrent is carried southward along the coasts of Ecuador and Peru, which are normally bathed by cold currents. The result is a wholesale destruction of animal life along the coast, including many guano birds that depend upon the sea for food. Permanent or semipermanent features of current systems that take a regular toll of life are found where the moving water carries the inhabitants into areas of less favorable living conditions. For example, Gulf Stream inhabitants ultimately perish as they are swept northward into regions where the temperature of the water is lowered by admixture of cold water or by cooling in higher latitudes. Larvae of neritic forms are frequently dispersed to offshore or other locations uninhabitable to the adult animals. Surface currents sometimes strew the shores with defunct bodies of normally oceanic or offshore forms such as the coelenterate Vellela or the pelagic snail Janthina.

Extent of the Marine Environment. That part of the earth which is capable of sustaining life, both plant and animal, is known as the biosphere. The biosphere is subdivided into three principal divisions or habitats known as biocycles. These are the terrestrial, the marine, and the fresh-water biocycles. Each has its characteristic types of ecological features and associations of plants and animals. A few animal species may at times migrate freely from one to another, as is witnessed especially by the salmon or the eel.

The oceans cover some 71 per cent of the earth's surface. Thus, the area of the oceans is about two and one half times the area of the land, but, when considering the space in which life might conceivably exist, account has to be taken of the relative vertical range provided by the two main environments, the terrestrial and the marine. On this basis it is estimated (Hesse, Allee, and Schmidt, 1937) that the marine environment actually provides about three hundred times the inhabitable space provided by the terrestrial and the fresh-water biocycles together; for, whereas the terrestrial environment provides space only in a shallow zone mainly at the immediate surface and to a depth of a few feet at the most, the marine habitat provides livable space for at least some form of life from the surface even to the abyssal depth of several miles. The fresh-water biocycle constitutes only a small fraction of the other two. The aerial portion of the globe is not properly considered a separate biocycle, since entrances into it by birds, insects, and so forth may be considered mainly as temporary journeys.


Owing to the difficulties attendant on the study of the oceans, the marine biocycle is the least known of all.

Classification of the Marine Environment

In order to facilitate a study of the marine environment and its inhabitants, the former may be conveniently divided broadly into primary and secondary biotic divisions based upon physical-chemical attributes or upon the nature of the biota. The boundaries between these biotic divisions, which are diagrammatically shown in fig. 67, are in some instances well defined, but more frequently there is a good deal of overlapping. Thus, although the primary divisions are definitely set off from each other on physical bases, and the typical subdivisions of these habitats can be clearly recognized both biotically and abiotically, yet there are no well-defined boundaries between them.


The main divisions of the marine environment.

The two primary divisions of the sea are the benthic and the pelagic. The former includes all of the ocean floor, while the latter includes the whole mass of water.

The Benthic Biotic Environment and Its Subdivisions. This division includes all of the bottom terrain from the wave-washed shore line at flood-tide level to the greatest deeps. It supports a characteristic type of life that not only lives upon but contributes to and markedly modifies the character of the bottom. Ekman (1935) discusses the boundaries of the vertical zones from a zoogeographic standpoint, and we follow mainly the scheme employed in his text.


The benthic division may be subdivided into two main systems—namely, the littoral and the deep-sea systems. The dividing line between these has been set at a depth of about 200 m on the arbitrary supposition that this represents the approximate depth of water at the outer edge of the continental shelf (p. 20), and, roughly, also the depth separating the lighted from the dark portion of the sea. The littoral system is subdivided into the eulittoral and the sublittoral zones. The deep-sea system is divided into an upper (archibenthic) and a lower (abyssal-benthic) zone. The limits of the benthic subdivisions are hard to define, and are variously placed by different authors because uniform boundaries that will fit all requirements cannot be drawn. For general biological studies, the different boundaries must be based on the peculiarities of the endemic plant and animal distribution and should follow the region of most distinct faunal and floral change. The biotic zones thus delineated will be characterized by a more or less clearly defined range of external ecological factors which have given character to the population.

The eulittoral zone extends from the high-tide level to a depth of about 40 to 60 m. The lower border is set roughly at the lowest limit at which the more abundant attached plants can grow. The sublittoral zone extends from this level to a depth of about 200 m, or the edge of the continental shelf. The dividing line between these subdivisions varies greatly between extremes, since it is determined by penetration of light sufficient for photosynthesis. It will be relatively shallow in the higher latitudes and deep in the lower latitudes. In the upper part of the eulittoral zone a relatively well-defined tidal or intertidal zone that is bounded by the high- and low-water extremes of the tide is recognized. Some authors confine the eulittoral zone to this narrow section and consider the sublittoral to begin at the low-tide level (cf. Gislen, 1930). The vertical range of the intertidal zone, though rather well defined for any given area, varies greatly in different sections of the world, for it is determined by the tidal range (see chapter XIV). In the upper reaches of the Bay of Fundy the zone may have a vertical range of over 15 m, while in the Gulf of Mexico it is less than 0.7 m, and in areas like the Mediterranean along the southwest coast of Italy the range is yet smaller, only 10 to 30 cm. On exposed coasts subjected to direct ocean waves and swells the upper range is somewhat extended to include a rather well-defined supratidal spray zone with a sparse population of especially resistant forms among which a few animals, such as the isopod Ligyda, appear to be in the process of becoming terrestrial in habit. Many species of animals are found only in the tidal zone and may be limited vertically in maximum distribution even to certain levels within the zone—for example, Ligyda and the gastropods Littorina scutalata, L. planaxis, Acmaea digitalis, and others found at Monterey Bay only above the 0.76-m tidal level (Hewatt, 1937). Thus, in the tidal zone

in which the range in external factors is greatest we find a more restricted vertical range of specific animals than is obvious anywhere else in the benthic region of the sea. Many motile animals, for example crustaceans and fishes, move regularly into the intertidal zone to feed during high tide, and the small pelagic fish known as grunion migrate into the zone during certain high spring tides to deposit their eggs in the sand.

The eulittoral zone gives rise to many biotopes, for it is greatly varied as to type of substratum—for example, rocky, sandy, or muddy—and also as to character of shore line and degree of exposure. The overlying water may be slightly or greatly reduced in salinity. These variations are direct, decisive features controlling the type and abundance of sessile littoral forms (cf. Shelford et al, 1935). The plentiful primary food in this zone is derived from both pelagic and attached plants.

Attempts to establish such zones as Fucus zone, Laminarian zone, and so on, based on the depths at which these plants are characteristically attached, has the disadvantage that the plants are very frequently absent along vast stretches of the coast, owing to unfavorable substratum or other ecological factors; nevertheless, such classification may be of useful local application.

Though the boundary between the sublittoral and the deep-sea systems is set at a depth of 200 m, Ekman's compilations based on the fauna indicate that in most regions the boundary may be located between 200 and 400 m. Light and temperature are important factors, and in high latitudes these factors operate together to shift the boundary into shallower water.

The upper division of the deep-sea system is called the archibenthic, a word introduced by Alexander Agassiz, but the term is unfortunate in that it implies the beginning of the benthos from this region. The zone is also called the continental deep-sea zone, but this gives rise to greater confusion, since the term “continental fauna” sometimes used must include also the littoral fauna unless specifically called continental-slope or deep-sea fauna. The archibenthic zone extends from the sublittoral to a depth between 800 and 1100 m.

The abyssal-benthic zone comprises all of the deep-sea benthic system below the archibenthic zone. It is a region of relatively uniform conditions. Temperatures are uniformly low, from 5° to −1°C, and solar light is wanting. There are no seasons, and hence the seasonal biological phenomena associated with the littoral zone are suppressed. Stagnant conditions do not prevail in the open ocean, however, for there is ample circulation to supply well-aerated water resulting from deep vertical movements in the high latitudes (p. 138). No plants are produced, and the extent to which autotrophic bacteria play a part in the manufacture of food is not known. The animals are carnivorous, feeding mainly upon organic detritus which in its initial organic state must have originated

in the plants of the surface waters. The abyssal zone, though not sharply marked off at its upper limits from the archibenthic zone, has its own characteristic population, as will be brought out in a following chapter.

The benthic environment from shore seaward to abyssal depths is covered, to a greater or less degree, by sedimentary deposits that may be classified as terrigenous deposits, organic or pelagic oozes, and red clay. A detailed discussion of the deposits will be found in chapter XX, and the nature of the distribution is shown in fig. 253. As far as the biology of benthic animals is concerned, the most important features of these oozes are their physical consistencies and the amount of digestible organic material they contain. Most deep-sea benthic forms are detritus eaters and mainly dependent, therefore, upon the rain of pelagic organisms that falls to the bottom. The production of pelagic food usually decreases markedly with increasing distance from the coast, and the amount reaching the bottom in areas of very deep water is further reduced by its disintegration while sinking. Hence, the littoral muds are most rich in food, and the red clay at great depths and far from shore is the poorest. This difference is reflected in the number of animals actually collected from different areas (cf. p. 806).

The Pelagic Environment and Its Subdivisions. The pelagic division includes all of the ocean waters covering the benthic division. Horizontally, the pelagic division is subdivided into an open-sea (oceanic) province, and an inshore (neritic) province.

Vertically, the oceanic province has an upper lighted zone and a lower dark zone with no well-marked boundary between the two. For convenience the boundary is arbitrarily set at 200 m, since this would correspond with the arbitrarily set depth for the edge of the continental shelf and at the same time place the littoral system and the neritic province in areas definitely within the lighted portion. Actually, light changes gradually in both quantity and quality from the very surface downward to depths where it is no longer detectable (p. 82), and this depth varies with latitude, season, amount of suspended material, living or dead, and therefore also with distance from shore. These variables of the pelagic environment are of profound importance to the population of the sea, as will be pointed out later.

The outstanding features of the oceanic province are the broad spatial expanses and the great ranges of depth. As distinguished from the neritic province the waters are as a rule very transparent, with little or no detritus of terrestrial origin. These waters are predominantly blue in color and support the blue surface fauna to be discussed more fully in chapter XVII. Although solar light penetrates relatively deeper than in inshore waters, the great depth of the water included in this province results in complete elimination of solar light in the deeper portion

of the province, and as a result only carnivores and detritus feeders can exist in the very deep layers.

The chemical composition of the offshore water is relatively stable. Salinity is uniformly high, with only small fluctuations in space and time (p. 123), and plant nutrients are frequently relatively low in the upper layer and only slowly replaced.

The vertical border separating the neritic province from the oceanic is set at the edge of the continental shelf; hence all water of depths shallower than 200 m would fall within the neritic province, which accordingly may extend far seaward in instances where the continental shelf is broad, as off the east coast of the United States, or be very narrow, as off the west coast of South America.

Although biologically and chemically the border between the oceanic and the neritic provinces is not strictly definable, yet as we approach the coast the plant and animal life takes on characteristics not found in the typically oceanic province where “blue-sea” forms prevail. The chemical constituents of the sea water in the neritic province are more variable than in the oceanic. Salinities are usually lower, sometimes markedly, and undergo seasonal or sporadic fluctuations such that many of the inhabitants are more or less euryhaline in nature—that is, able to endure wide ranges of salinity. River water may bring in nutrients and may also exert a stabilizing influence on the turbulent motion, being at times, therefore, instrumental in initiating plant growth in the upper layers (p. 789). Plant nutrients, nitrates, phosphorus, and so on are more readily available in the shallower inshore water because of the greater possibility of return by vertical currents after they have been regenerated from the disintegrating organisms on the bottom or in the deeper water (chapter VII). This factor is of the utmost importance to production of diatoms, foremost of the primary food of the sea. Therefore, per unit area of the sea, the neritic province is far more productive than the oceanic province and is consequently the region of greatest importance to marine life in general. Here fish of greatest economic importance are taken, not only because of greater availability, but also because it is their natural habitat.

Other Biotic Units. The above classification of the marine environments is based mainly on broad geographical, physical, chemical, and biological characteristics that circumscribe more or less clearly the separate zones. Within each of these extensive zones we observe many and varied sets of ecological conditions resulting from differences in substratum, proximity to shore, depth and chemical-physical condition of the water, and so forth.

The primary “topographic” unit used in ecological classification of the environment is the biotope, or niche, which is defined as “an area of which the principal habitat conditions and the living forms which are

adapted to them are uniform” (Hesse, Allee, and Schmidt, 1937). Since in any given type of biotope the habitat conditions make specific demands on the inhabitants, it follows that an analogous development of the inhabitants is frequently reflected in the population, and those not fitted for the habitat are eliminated from it. Obviously, some organisms are not so narrowly bound to the biotope as are others of more specialized nature. Thus, within a biotope may be found some generalized forms such as certain cephalopods and fishes that wander more or less freely from one type of biotope to another. The more specialized a biotope becomes with respect to living conditions, the more uniform will the inhabitants become, so that only a few species with large numbers of individuals may exist. The smaller habitat anomalies found within the biotope are called facies. The number of organisms that can live in any given biotope may in special cases be determined by available suitable space, but more frequently it will depend upon the food supply that may be produced within the biotope or be carried to it from outside by currents. The community of forms in a biotope is called a biocoenosis.

Biotopes having certain characteristics in common—for example, proximity to the coast or estuarine locality—are united into larger divisions known as biochores.

General Character of Populations of the Primary Biotic Divisions

Under the previous headings we have dealt with the classification of the marine environment. For purposes of future discussion it is desirable at this point to outline briefly a broad, highly practical classification of the marine population inhabiting the above primary biotic divisions, a classification based not on natural phylogenetic or taxonomic relationships, as given on p. 282, but rather on an artificial basis, grouping heterogeneous assortments of organisms depending upon common habits of locomotion and mode of life and upon common ecological distribution.

On these grounds the population of the sea may be divided into three large groups—namely, the benthos, nekton, and plankton, the first belonging to the benthic region and the other two to the pelagic region.

In the benthos (Gr., deep or deep-sea) are included the sessile, creeping, and burrowing organisms found on the bottom of the sea. Representatives of the group extend from the high-tide level down into the abyssal depths. The benthos comprises (1) sessile animals, such as the sponges, barnacles, mussels, oysters, crinoids, corals, hydroids, bryozoa, some of the worms, all of the seaweeds and eel grasses, and many of the diatoms, (2) creeping forms, such as crabs, lobsters, certain copepods, amphipods, and many other crustacea, many protozoa, snails, and some bivalves and fishes, and (3) burrowing forms, including most of the clams and worms, some crustacea, and echinoderms.


The nekton (Gr., swimming) is composed of swimming animals found in the pelagic division. In this group are included most of the adult squids, fishes, and whales—namely, all of the marine animals that are able to migrate freely over considerable distances. Obviously, there are no plants in this general group.

In the plankton (Gr., wanderer) is included all of the floating or drifting life of the pelagic division of the sea. The organisms, both plant and animal, of this division are usually microscopic or relatively small; they float more or less passively with the currents and are therefore at the mercy of prevailing water movements. Many of the animals are able to make some progress in swimming, although their organs of locomotion are relatively weak and ineffective. The plankton is divided into two main divisions, the phytoplankton and the zooplankton. The former comprises all of the floating plants, such as diatoms, dinoflagellates, coccolithophores, and sargassum weeds. In the zooplankton are included (1) myriads of animals that live permanently in a floating state, and (2) countless numbers of helpless larvae and eggs of the animal benthos and nekton. Since the plankton and nekton occupy the same biotic realm and are part of the same community, it is necessary always to remember that the distinction is one based primarily on relative size and speed of swimming, and does not signify a divergence of ecological relationship.

Each of these three ecological groups will be more fully discussed in later chapters.

Development of life in the Sea

Let us review briefly the observations that indicate the relative antiquity of the marine environment as a biological realm. It is not possible to know when life arose in the sea, but the close similarity of the chemical composition of body fluids and sea water has led to the supposition that the sea was already saline at that early time and that, because of the intimacy of primitive organisms with the fluid environment, the elements present entered into the fundamental composition and mode of metabolism of the primitive organisms and are maintained in present-day forms with certain modifications in the proportions of the principal ions, especially magnesium (table 55). These interesting relationships have led to much speculation relative to the development of organisms and the chemical composition of primitive seas, but we cannot enter further upon that phase of the action of the environment. Pearse (1936) has given some reviews and listed literature pertaining to these questions and to the theory of migration of animals from sea to land.

The part played by the sea in the distribution and maintenance of present-day life upon our globe is a vital one. The sea itself is abundantly populated, and no life could exist on land were it not for the perpetual

water cycle of evaporation, precipitation, and drainage between sea and land. Only in the sea would it be possible to approach any degree of self-sufficiency as a biological realm, and historically the sea has acted a principal role in the development of animal life.

That the sea is the original environment of animal life is strongly indicated by certain facts that point to the greater age of marine life as compared to terrestrial and fresh-water faunas to which it has seemingly given rise. Evidence pointing to a greater age of marine fauna over the terrestrial and fresh-water faunas is mainly along four lines: (1) general composition of present-day faunas, (2) similarity in the chemical composition of body fluids and sea water, (3) life histories, and (4) paleontological relationships.

(1) The whole animal kingdom is divided into a number of primary divisions, each known as a phylum. Each phylum is composed of animals having certain fundamental morphological similarities not possessed by any animals of other phyla. Thus, a natural, as opposed to artificial, relationship is indicated. Each phylum is then divided into natural but more restricted groups known as classes, and these in turn are followed by other yet lower divisions in the following manner:


Species are formed of individuals, and the morphological features by which each species is characterized are less fundamental and presumably of more recent origin than those characterizing the genera. Similarly, the generic structures are less fundamental than those of families, and so on to the highest division, which is based on structures of great antiquity.

A review of all the higher or major divisions—namely, the phyla and classes of animal life—reveals the striking preponderance of marine groups. All of the seventeen phyla (using the taxonomic ranking of H. S. Pratt, 1935, in Manual of Invertebrate Animals) are represented in the sea, and most, if not all, are believed to have originated there. The following five are exclusively marine: Ctenophora, Echinodermata, Phoronidea, Brachiopoda, Chaetognatha. Some authors recognize fewer than seventeen phyla, but this has only the effect of increasing the preponderance of purely marine classes.

Of the forty-seven classes (where only subphyla were given under phyla, they are here rated as classes) of invertebrates as given by Pratt, twenty-one, or 43.7 per cent, are exclusively marine, and only three, or

6.2 per cent, are exclusively nonmarine. Of the subphylum Vertebrata, only members of the class Amphibia are nonmarine, while the other four classes share members in both marine and nonmarine environments. The fishes are predominantly marine, while the reptiles, birds, and mammals are predominantly terrestrial. The amphibians represent the highest nonmarine group.

These divisions demonstrate the astonishing variety of marine animals, as far as the major phylogenetic groups are concerned. However, the terrestrial environment harbors the greatest number of species, mainly owing to the large number of species of one restricted group, the insects, which are almost totally absent from the sea. The presence in the sea of so many major groups, many of which are restricted to the sea, indicates the great tendency on the part of the marine environment to preserve the groups that have once become evolved.

It should be noted also that, in addition to the remarkable diversity of marine life in the ocean, there is a conspicuous primitive element, as judged by simplicity of structure, in the groups represented. In the sea there is a more complete developmental series of animal life than exists anywhere else, because of which, and also because of the natural and intimate relationships of the organisms to the sea-water medium, the studies issuing from the marine biological laboratories have contributed vastly to information on biological problems dealing with development and maintenance of life.

The relative uniformity of the marine environment has been instrumental not only in preserving the diversity of forms but also in retaining a generally more primitive character as compared with terrestrial and fresh-water animals. It is true that in the sea we do find associated with the lower forms a number of highly developed animals that must be considered marine because of their dependence on the sea. These are the seals, whales, certain reptiles, fishes, and birds. All of these groups, however, have had a large part of their racial development in the terrestrial and fresh-water habitat. They have more recently reverted to the sea and have only secondarily become adapted to it. The teleost fishes, which are believed to have evolved to their present status in fresh water, were originally derived from marine stock.

(2) The relation of body fluids to sea water has already been discussed (p. 269).

(3) A study of the life histories of invertebrates suggests the antiquity of marine life. During the early history of the individuals of some animal groups the larval stages are markedly different in structure and habit from the mature phase. The larval stages, which sometimes resemble the mature stages of other groups or only the larvae of other groups, are thought to reflect a structural similarity to ancestral stock. Whether or not this is a real recapitulation of racial history or only an

expression of individual larval adaptation to a common environment is important in seeking an understanding of the similarity. Whatever the truth may be, it is well known that most marine invertebrates pass through an early stage during which the larvae in no way structurally suggest their parentage, but may even have striking fundamental similarity to existing larvae of other groups. From this it has been possible to establish types of larvae—for example, the trochophore of the Annelida and Mollusca, and the nauplius of the crustacean groups (fig. 80, p. 321).

There is a tendency for some aggressive animal groups to desert the sea for fresh-water or land habitats. This is shown by the crustaceans, among which there are forms such as the prawn, Eriocheir, which enters fresh water at a young stage but when mature returns to the sea to spawn. The land crabs, Cardisoma, Gecarcinua, and so forth, also go through a free-swimming larval stage in sea water.

(4) It is well known that animal fossils occurring in the oldest known fossiliferous rocks of the earth's crust are mainly marine forms.

Marine animals were abundant and became fossilized in the Cambrian period (500 million years ago), when certain portions of the land now above sea level formed a part of the sea bottom along the coasts of ancient seas. Several invertebrate phyla were already developed, and such forms as trilobites and brachiopods were particularly abundant.

The chief roles of the marine and terrestrial environments in the development of life may be summarized by saying that the great part played by the former is chiefly in the development and maintenance of a wide diversity of lower forms, while in the latter the influence of the more rigorous habitats has produced less diversity of form but a higher type of complexity.

The area where these two great environments meet, the intertidal zone, is in an intermediate position and subject to rapid and marked vicissitudes, and it is from here that much of the migration to land is supposed to have taken place.


Bethe, A.1929. “Ionendurchlassigkeit der Körperfläche von wirbellosen Thieren des Meeres als Ursache der Giftigkeit von Seewasser abnormer Zusammensetzung” . Pflugers Arch., 221, p. 344–362, 1929.

Dakin, W. J.1935. “The aquatic animal and its environment” . Linnean Soc. New South Wales, Proc., v. 60, pts. 1, 2, p. viii–xxxii, 1935.

Ekman, Sven. 1935. “Tiergeographie des Meeres” . Akad. Verlagsgesellsch., Leipzig. 542 pp., 1935.

Gislen, T.1930. Epibiosis of Gullmar Fjord. II. Kristinebergs Zool. Sta. 1877 to 1927, No. 4, p. 1–380, 1930.

Hesse, Richard, W. C. Allee, and K. P. Schmidt1937. Ecological animal geography. An authorized, rewritten edition based on “Tiergeographie auf oekologischer Grundlage,” by Richard Hesse. John Wiley & Sons. New York. 597 pp., 1937.


Hewatt, Willis G.1937. “Ecological studies on selected marine intertidal communities of Monterey Bay, California” . Amer. Midland Naturalist, v. 18, p. 161–206, 1937.

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Pearse, A. S.1936. “The migrations of animals from sea to land” . Durham, N. C., Duke Univ. Press, 176 pp., 1936.

Pratt, Henry S.1935. “A manual of the common invertebrate animals exclusive of insects” . Revised. Philadelphia. Blakiston, 854 pp., 1935.

Shelford, V. E., et al.1935. “Some marine biotic communities of the Pacific Coast of North America. Pt. 1. General survey of the communities” . Ecol. Monographs, v. 5, p. 250–332, 1935.


IX. Populations of the Sea


In the sea, as on land, the plants are the real producers—that is, the organisms that are capable of elaborating complex organic substances from the simple inorganic compounds dissolved in the water. Without marine plants as synthesizers of primary food, development of marine animal life would be impossible beyond a negligible quantity that might be supported alongshore and in estuaries where particulate organic material of terrestrial origin would find its way into the sea.

A notable feature of marine vegetation is its dearth of variety when compared to the multiplicity of forms characterizing the terrestrial vegetation. Also, the types of plants most important in the production of primary food in the sea are in striking contrast to those constituting the chief synthesizers on land. This marked difference is readily explained, as we shall see, since it is dependent upon the radically different demands made on the plants by the marine environment. The poverty of plant variety in the sea is also in striking contrast to the abundant diversity of marine animal life. It may truly be said that the animal kingdom belongs mainly to the sea, while the terrestrial environment fosters the plants, although the most primitive of the plant groups, the algae, are wonderfully developed in the sea.

Light is of prime importance to all photosynthetic plants, and the possibility for attachment to the substratum is of secondary importance. More will be said about this later, but we must point out here that only in a very small portion of the sea are the two factors, light and suitable substratum for attachment, at the same time operative. This small portion of the sea wherein there may be sufficient light penetration to support attached plants—that is, the eulittoral zone—constitutes about 2 per cent of the sea floor.

Anyone frequenting the seashore is familiar with the covering of brown rockweed, Fucus, the green sea lettuce, Ulva, and a number of other low-growing plants that carpet the rocks in the intertidal zone. These or yet other relatively low-growing or encrusting plants may extend to varying depth below low tide if a suitable substratum for attachment is

available. Bottoms of mud or sand can furnish attachment only where scattered larger rocks are exposed above the smaller shifting particles, since the algae do not possess true roots for anchorage in the soil. A few types of algae—for example, species of Caulerpa—are able to bind themselves to sand in such a manner that the many slender, long branches of the holdfast may enclose a ballast of sand. Shingle beaches may give anchorage to the plants only to the extent that individual rocks are not regularly shifted by the action of currents and waves. Many of the smaller algae are epiphytic, growing on other plants, or even epizoic, growing on animals, but in general most of the attached forms may be considered lithophytic.

The large kelps, such as Nereocystis, Pelagophycus, and Macrocystis, are found typically on rocky reefs some distance from the intertidal zone. Growth may occur on shoal reefs or rocks miles from shore, but the destructive mechanical effect of breaking waves and swells usually prevents any growth of these long-stiped forms in the immediate vicinity of exposed shores or rocks. Hence, the large kelps characteristically form in bands or patches some distance from shore where there is active circulation of water and yet where the danger of abrasion is reduced.

It was pointed out that only a small per cent of the sea floor may be considered to have sufficient light to support attached plants. Although this area may have enough light, it is vastly reduced as a suitable area for attachment of larger plants because of the great coastal stretches of mud, sand, shingle, or other unfavorable features. Therefore the bulk of the material produced by the attached marine plants is relatively small and can support only a small portion of the animal life actually present throughout the vast marine habitat; nevertheless, in more restricted areas along the coasts, attached plants—for example, the eel grasses—may be the chief producers. As a result of this restricted production by the benthic, or attached, plants, the primary food production becomes mainly a function of the unattached floating plants, notably, the diatoms and dinoflagellates, which, though microscopic in size, occur in vast, incalculable numbers.

Accordingly, our study of plant production must be concerned mainly with these floating forms. The means and adjustments by which this extensive community of floating plants—that is, the phytoplankton—is maintained and is related to other forms of life will be dealt with in subsequent chapters. First, however, in order to have a more complete understanding of the whole biological “setup” of the sea, it will be necessary as a point of departure to make a brief review of the various groups that are important to the economy of the sea as a whole.

The entire plant kingdom is divided into four primary divisions: the Thallophyta, Bryophyta, Pteredophyta, and Spermatophyta. Only the first and last of these are represented in the sea.


These primary divisions are each again divided and subdivided into many smaller secondary divisions which are indispensable to the specialist in marine botany, but for our purpose it will suffice to mention these smaller divisions only when they are important to marine economy, or when they have obtained rather widespread inclusion in more or less general literature dealing with the sea.

Only an abridged classification can be given. For a more complete treatment of the systematics the reader can refer to numerous good texts on botany or to publications dealing specifically with the group in which he is interested. A few of these publications are included in the bibliography, and yet others can be traced from those works included.


Nearly all of the marine plants fall into this botanical division, which is made up of primitive plants in which the body shows little or no differentiation of vegetative organs—that is, no true root, stem, or leaf. Important among these thallus plants are the marine algae and the marine fungi, especially the bacteria. Since bacteria constitute the subject of a more specialized study of the sea, they will be dealt with under a special heading in chapter XVIII.

Most algae are beautifully colored, and sometimes also iridescent. The pigments of the chromatophores intercept solar energy, which is used in the synthesis of organic compounds. The type of pigment or pigment combination occurring in the algae as color manifestations has led to the names commonly used for the classes:

  1. Blue-green algae (Myxophyceae)

  2. Green algae (Chlorophyceae)

  3. Brown algae (Phaeophyceae)

  4. Red algae (Rhodophyceae)

Yellow-green algae (a heterogeneous group variously classified by different authors)

In general, the colors are characteristic of the classes, but other characteristics associated with cell structure and life history are more fundamental in distinguishing the five groups. Each group has a considerable variation in general morphology, some features of which will be pointed out in a review of the classes. The first four, with the exception of some blue-greens, are attached plants, while the yellow-greens are characteristically floating, or planktonic, forms.

Blue-Green Algae (Myxophyceae)

This class contains only small, poorly organized plants, some consisting of only a single cell, while others are multicellular. The blue color of these plants is due to a water-soluble accessory pigment, phycocyanin. In certain inland waters, it has been reported that upon the

death of large masses of a blue-green alga (Anabaena), the water-soluble pigment may impart a deep-blue color to the waters. The Red Sea owes its name to a free-floating form, Trichodesmium erythraeum (fig. 68a), which has a red accessory pigment and is responsible for the red color sometimes observed in the surface waters. Thus, “blue-green algae” may be red. The cell walls of plants of this group consist usually of chitin, instead of the cellulose so characteristic of other plants; therefore, in a small measure, they supplement the enormous quantities of chitin produced in the sea, especially by the crustacea. Some Myxophyceae are endophytic; that is, they live within the bodies of other plants in an association known as symbiosis. A marine species, Richelia intracellularis, may be found within the cells of the diatom Rhizosolenia.


Characteristic types of multicellular marine algae. a, Trichodesmium; b, Fucus; c, Alaria; d, Ulva; e, Ectocarpus; f, Sargassum; g, Rhodymenia; h, Polysiphonia; i, Cytosiphon; j, Lithothamnion.


Methods of Reproduction. Reproduction in this group is by asexual fission. This, the most simple method of propagation, consists of single individuals dividing to form two of lesser size, which in turn again divide after growth. In instances involving blue-green algae that form chains of cells, the chains divide into smaller sections known as hormogonia. Fission of the cells in the hormogonia again increases the length of the filaments.

Distribution. The Myxophyceae are of less general importance in the oceans than are the following algal groups. They are widely distributed in fresh and brackish water. In the sea they are most often found in the warmer waters, where they may cause the phenomenon of sliming. A brackish-water form, Nodularia spumigena, native to calm fjords of the north, may at times cause extensive sliming of the waters. In the Gulf of Bothnia, sliming due to this or similar forms may assume considerable proportions.

Green Algae (Chlorophyceae)

As the name indicates, the algae of this class are green in color. The pigments of the chloroplasts include the two types of chlorophyll, a and b, and the various carotinoids. The yellow and orange of the latter pigments are masked by the abundance of the green chlorophyll. In contrast to the chitinous cell wall of the blue-greens, these plants produce walls that are largely cellulose—a carbohydrate as opposed to the nitrogenous product, chitin. Some green algae of the sea—for example, Halimeda of the Siphonales—become incrusted with calcium carbonate, and thus may contribute materially in some places to the formation of lime deposits in warmer seas. The joints of the plant remain uncalcified, and thus allow flexibility in the moving water.

There is great diversity in the morphological features of this class. Common forms are filamentous with septa (Urospora) or without septa (Codium), tubular (Enteromorpha), and sheet-like (Ulva, or sea lettuce) (fig. 68d).

Methods of Reproduction. Common methods of reproduction may be illustrated by the habit of the cosmopolitan Ulva. In sexual reproduction the contents of any of the ordinary cells of the flat two-layered plant may form biciliated bodies called gametes which, upon escaping into the water, unite in pairs and by cellular division grow to form the new plant, known as the sporophyte, but usually passing first through a filamentous stage. Reproduction may also be asexual, in which case any of the common cells of the sporophyte plant may form microscopic quadriciliate zoospores (spores are simple reproductive cells which differ from seeds mainly in that they do not contain any ready-made embryo plant). These zoospores, upon being discharged, grow directly into gametophytes, the plants that produce the gametes.

This process is known as alternation of generations, and associated with it are important cytological changes. In the case of Ulva, however, the macroscopic features (the sporophyte and gametophyte plants) are indistinguishable—that is, they are isomorphic or homologous. Asexual reproduction may also occur by fragmentation, thus forming unattached plants.

During the period of reproduction, large swarms of gametes and zoospores may be released, leaving the parent plant colorless and forming a green “bloom” on the waters of quiet bays. For many filter-feeding animals, the floating microscopic reproductive products of these and other algae form a source of food that must not be overlooked in a study of food of littoral animals. In bays, also, these swimming stages of algae, as well as algal slime, contribute to primary film formation that leads to an eventual fouling growth on ships and other submerged structures.

Distribution of Green Algae in the Sea. The green algae are found mainly in the upper littoral zone, especially in the lower half of the tidal zone, and in the immediate subtidal region down to a depth of 10m or more, and therefore in a relatively well-lighted habitat. It is with the green algae that the fresh-water algae are most closely related.

In geographic distribution, green algae are found most abundantly in the warmer seas. Algologists have remarked on the relative scarcity and dwarfed development of the Chlorophyceae in the Arctic Sea.

Brown Algae (Phaeophyceae)

Brown algae belong almost entirely to the sea, only a very few occurring in fresh water. Here are included the conspicuous brown seaweeds, many of which grow to notably large size. The pigments of this class include green chlorophyll, which is masked by the yellow and brown pigments, xanthophyll, carotin, and fucoxanthin.

Plants of this class of algae form the conspicuous offshore growths popularly known as “kelp beds.” They are the giants among the seaweeds, and form the marine forests among whose waving stipes and fronds myriads of neritic fish obtain their food and seek shelter from their aquatic enemies. These, also, are the kelps commonly harvested in many places for the commercial products they yield.

The brown algae possess a great range of size and structure. There are minute, delicate, filamentous branching plants (Ectocarpus, fig. 68e); coarse, hollow, sausage-like chains a foot or more in length (Scytosiphon, fig. 68i); short-stalked forms with broad thalli (Laminaria, Costaria, and Alaria, fig. 68c, some of which become nearly 2 m broad); many branched forms (Fucus, Egregia); and long-stalked giants of the Pacific with long leathery fronds (Macrocystis, Nereocystis, Pelagophycus).

In structure the brown algae are the most advanced of all thallophytes. If we refer only to the more superficial details, Nereocystis

(fig. 69) will serve to illustrate the essential features of the larger typical brown seaweeds and to give a basis for interpretation of structural features of other groups as well.

Nereocystis may attain a length of 35 m or more. The plant is anchored to a hard substratum by means of a profusely branched structure known as the holdfast, but there are no true roots. From the holdfast extends the long cylindrical stipe, which is hollow through most of its length and ends distally in a large hollow bulb. This bulb, like the stipe, is filled with gas, giving buoyancy to the plant. Ribbon-like fronds or laminae issue from the distal end of the bulb.


The gross structure and life cycle of Nereocystis. a, sporophyte plant; b, swimming zoospores; c, male and d, female gametophyte plants; e, young sporophyte.

The hollow bulb and stipe maintain the upper portion of the plant near the surface, exposing the fronds to favorable light conditions. In common with other large algae, the parts are tough, flexible, and slippery in order to withstand with least resistance the effect of frequently violent storm waves and strong currents.

Methods of Reproduction. The life cycle of the brown algae includes various types of alternation of generations. Commonly, in the Laminariales, which includes the large kelps, there is an alternation of generations that may be illustrated by the cycle shown in Nereocystis (fig. 69). Here the large, conspicuous sporophyte plant produces a series of sori, or “fruiting areas,” appearing as dark brown patches running longitudinally along the whole length of the fronds. Beginning at the distal end of the frond, these patches are detached at maturity, leaving a broad gap (3 to 10 cm) in the frond. From the mature sori, innumerable ciliated zoospores escape and, upon reaching a suitable substratum, grow to small filamentous plants, the inconspicuous gametophyte stage. Thus the alternation of generations in this form is

heteromorphic. A few brown algae—for example, Dictyotales—show isomorphic alternation of generations, as discussed for the green alga Ulva. It is of significance to note that a conservation of zoospores may perhaps be effected in Nereocystis by the habit of shedding the complete, mature sorus, which, upon sinking to the bottom of the kelp bed, is more likely to find a favorable substratum. Thus, the zoospores when released are concentrated at or near the bottom, rather than widely dispersed by the currents as would have been the case if they had been released from the sporangia of the sori at the surface of the water. According to Hartge (1928) the zoospores germinate in twenty-four hours. The resulting gametophyte plants are either male or female, and, upon fertilization of the egg, growth of the sporophyte is initiated.

In the brown algal group, Fucales, to which Fucus and Sargassum (fig. 68f) belong, the main plant is a sporophyte, but, within the thousands of tiny cup-like conceptacles forming the bladders, gametes are formed like spores. These unite after being discharged free in the water. Thus the alternation of generations is evident only cytologically. In connection with the “spawning” of Fucus, it is interesting to note that it is rhythmic with the tide, taking place after a period of exposure at low tide.

Distribution. The brown algae reach their maximum development in cooler waters, and are therefore typical of the rocky coast of higher latitudes. Sargassum and others of the Fucales are characteristic, however, of tropical or subtropical regions. Tilden (1935) is of the opinion that the Laminariales arose in the North Pacific, while the Fucales had their origin in the South Pacific. Several species or varieties of Sargassum, or “gulfweed,” are found in large quantities in the Sargasso Sea, whence they have drifted and multiplied after being torn loose from coastal areas. They are kept afloat by air bladders and grow vegetatively, propagating by fragmentation, but apparently do not form fruiting bodies. The drifting masses form a characteristic environment with associations including other algal and animal forms of littoral type.

The vertical distribution of brown algae shows many low-growing forms, especially the Fucales, in the rocky intertidal zone. Near the lowest tide level the medium-sized forms with leathery fronds and short limber stipes begin to prevail, and they increase markedly in the next 15 to 20 m of depth, finally diminishing and disappearing below the eulittoral zone.

Intermingled with these short-stiped algae are the giant long-stiped kelps that usually grow most abundantly some distance from the shore and extend to depths of 30 m or more. Macrocystis, one of the giant kelps of the Pacific, is said to reach to the surface from a depth of 80 m off the coast of Chile (Hesse, Allee, and Schmidt, 1937), but in the North Pacific it has its most abundant growth in water of about 15 m. Kelps of this genus are said to be absent from strictly tropical waters and, as a

group, may be found in waters with temperatures from −2° to nearly 25°C. A few species range through all of the degrees of temperature, but most kelps are confined to narrower limits (Setchell, 1912).

Mention should also be made of epiphytic forms such as the filamentous Ectocarpus (fig. 68e), which prefers to be attached to other algae growing at various depths.

Red Algae (Rhodophyceae)

Nearly all of the red algae are marine. From the standpoint of color, they are the most striking of all the marine algae, some of them being also highly iridescent. Many of the delicate forms are among the most beautiful macroscopic objects of the sea. The order Gelidiaceae ranks first in importance commercially since certain of its members form the main source of agar.

The pigments of the chromatophores include the usual chlorophylls together with xanthophyll, carotin, and, in addition, the red phycoerythrin and sometimes phycocyanin. The plants may appear red, purple, violet, or, to some degree, brown or green. The deeper-growing species are the more purely red, a fact which is perhaps associated with their ability to synthesize more efficiently in the subdued light of greater depths than are the shallow-water types (Gail, 1922).

Though usually small in size, the red algae show a diversity of form much greater than the brown, and they are also more numerous. All are multicellular, the simplest being filamentous branching forms like Polysiphonia (fig. 68h), which, together with other filamentous algae, are commonly called “sea moss.” The larger flat types may be illustrated by Rhodymenia (fig. 68g), in which the broad frond may attain a considerable length. However, the maximum length of the larger red algae is only about 1 to 2 m.

Methods of Reproduction. The life cycle of some species is very complicated and cannot be amply discussed here. The reader is referred to the works of Kylin and other texts for a more complete treatment. In the higher types there is a regular morphological alternation of generations in which the sporophyte and gametophyte may superficially appear similar. Polysiphonia is commonly used to illustrate the life cycle of red algae. Here three types of plants are produced—namely, a male and a female gametophyte and an asexual tetrasporic plant. The last arises from the carpospores, which occur on the female plant. The carpospores are the products of union of male and female gametes. Upon germination, the tetraspores of the asexual plant give rise, in turn, to the sexual plants.

One of the most remarkable features of reproduction in red algae however, is the complete absence of any ciliated or flagellated swimming spores or gametes. This feature is a notable departure from the rule

followed in reproduction of organisms occurring in a water medium. It makes the dispersal and ultimate contact of reproductive cells dependent upon currents and hence wholly a matter of chance.

Distribution. The Rhodophyceae are widely distributed geographically, but are most abundant in temperate seas. Their vertical distribution indicates that they prefer to grow in subdued light. A few species may be found in the intertidal zone, but the most luxuriant growth is subtidal. They may occur in abundance in depths less favorable to most of the green and the brown algae, and in the Mediterranean they have been reported from depths of 130 m. Thus, from shallow to deep water the general vertical distribution of the algal groups discussed is successively the green, the brown, and the red, with a wide degree of overlapping.

It should be mentioned here that certain red algae (the Nullipores) play an important role in calcium carbonate precipitation in the sea. They have contributed, and still do contribute, greatly to geological formations. Among these are, especially, the coralline algae, of which Lithothamnion (fig. 68i) is a typical example. They are distributed from lat. 73°5′ S to 79°56′ N (Tilden, 1935) and can be observed as copious encrustations on rocks and shells in the littoral zone of every exposed shore.

Yellow-Green Algae

There is considerable disagreement as to the proper grouping and status of divisions within this heterogeneous assemblage of organisms, some of which, as indicated below, are animal in nature. As a matter of convenience in discussing the more important marine members, we shall here employ only names of more or less familiar usage in biology and oceanography. Many of the members included are classified as animals in zoological texts, but in consideration of their holophytic nature (faculties of photosynthesis) it is most convenient for oceanographic studies to include them a priori among the producers. For more detailed treatment of the systematics of the various divisions, the reader is referred to Fritsch (1935) and the relevant works included in the discussions under the separate groups.

In contrast to the algae previously discussed, the members of this assemblage of plants and plantlike animals are primarily floating forms and will be taken up in the order of their importance in the economy of the sea.

Diatoms. The plants here included are all microscopic in size, the larger species viewed individually appearing only as tiny points. Some earlier authors of marine botany included them with the brown algae. A comprehensive treatment of the group is given by Hustedt (1930). In structure they are unicellular, but individuals may form chains or groups of various types. Examples of types representing the common genera are

given in fig. 70. A characteristic feature of diatoms is the shell, or frustule, which is composed of translucent silica, and the variety of form and sculpturing shown by striae, points, and pits is truly remarkable. These shells are of considerable importance in siliceous sediments and have formed great fossil deposits known as diatomaceous earth.


Characteristic types of diatoms. a, Corethron; b, Nitzschia closterium; c, Planktoniella; e, Coscinodiscus; f, Fragilaria; g, Chaetoceros; h, Thalassiosira; i, Asterionella; j, Biddulphia; k, Ditylum, l, Thalassiothrix; m, Navicula; n, o, Rhizosolenia semispina, summer and winter forms.

Since these plants as a group may be considered the most important in the economy of the sea, it is imperative that we treat them in considerable

detail in order to facilitate understanding of their mode of maintenance, numerical fluctuations, and requirements in the sea as set forth in the following chapters, in which organisms are considered in the light of ecological factors.

The shell structure of diatoms (fig. 71) may be likened to a box with a telescoping lid, because it consists of two nearly equal halves fitted one over the other. The pieces corresponding to the top and bottom of the box are known as the valves, and these are each joined by connecting bands that overlap and together form the girdle. The larger half of the shell is known as the epitheca, and the smaller half, which fits into it, as the hypotheca. The protoplasm lies wholly within the shell, but for exchange of metabolic products it is exposed by a slit (raphae) in the valve of some types and by small pores in others.


The gross structure of a simple diatom (Coscinodiscus). a, valvular view; b, girdle-view section of cell wall.


Reproduction in diatoms. a,b, cell division; c, diminution of size resulting from cell division in three generations.

Diatoms may possess only one or many chromatophores, which may vary in color from yellow to olive-green or brown. Authorities are in poor agreement as to the nature of the pigments present, but there is some indication that the common pigments are masked by the accessory brown pigment diatomin, which may be identical with fucoxanthin of the brown algae. An important product of assimilation is an oil that is frequently visible as droplets within the diatom.

Methods of Reproduction. The most common method of propagation among the diatoms is by simple cell division (fig. 72a). This method has a far-reaching effect on the population in two distinct ways. First, it is conducive to a rapid production of enormous numbers when

conditions for growth are favorable (p. 767). Second, the maximum size attainable by individuals in a portion of the population is constantly being reduced by each successive division. This reduction follows from the fact that as a result of division of the protoplast (living portion of the diatom) one of the new daughter cells retains the larger of the valves (epitheca), whereas the other daughter retains the smaller valve (hypotheca). During the process of division, new complementary valves are laid down, but in such a manner that they fit into the old valves of the parent; that is, they become hypothecae, and the hypotheca of the parent diatom must now be looked upon as the epitheca of the smaller of the two daughter diatoms formed. Thus, through many successive generations of uninterrupted cell division a marked reduction in width must occur in many individuals. This change is schematically illustrated for three generations in fig. 72b. It appears that this diminution of size can go on only to a certain point, when a return to maximum size must be accomplished by the formation of auxospores (fig. 73a). If auxospores are not formed at a certain minimum size, the decrease in size continues, with ultimate abnormalities and death. In auxospore formation the contents of the diminutive, rigid, siliceous shell escape from the parted valves inclosed in a distensible pectin membrane. It is then possible for them to grow to larger size with the formation of new full-sized valves. A number of variations occur in the method of auxospore formation, dependent upon the group. For example, several auxospores may result, and these may fuse with yet others. In general, however, some type of rejuvenation seems to take place. Auxospore formation has been shown in various species, but in nature the spores are found only in small numbers.


Reproduction in diatoms. a, auxospore formation in Thalassiosira aestivalis (after Gran and Angst); b, increase in cell size following auxospore formation in Melosira nummuloides (after Fritsch); c, resting spores in mother cells, Chaetoceros vanhurckii; d, resting spore of Chaetoceros diatema; e, resting spore of Chaetoceros radicans; f, microspores in Ditylum; g, microspores in Chaetoceros didymus (after Gran and Angst).


Diatoms may also produce what are known as microspores (fig. 73b). These were early observed by Murray, Gran, and others. They consist of small protoplasmic spheres occupying the shell, and may escape as biciliated spores. The significance of these bodies is not fully known.

Resting spores of characteristic structure (fig. 73c) are also formed in most pelagic neritic species, especially of the centric types, by the cell contents becoming condensed and surrounded by a heavy, siliceous wall. They may be produced at the initial appearance of unfavorable living conditions, and may drift for some time within the old frustule or sink to the bottom to survive the unfavorable seasons of inadequate nutrients, cold, or of varying salinity so characteristic of many coastal areas. Gran (1912) has reported them from Arctic collections in which they were enclosed in ice.

Winter and summer forms of oceanic diatom species have been reported. These are cases of marked dimorphism in which the coarse winter forms have been looked upon as a means of survival from one favorable season to another. However, the dimorphism may be only an adjustment to changes of viscosity inherent with seasonal temperature changes.

Many diatoms grow normally on the bottom in the littoral zone, where they may or may not be attached by stalks or glide freely over the bottom. These benthic forms produce the heavily shelled types with most exquisite designs. Diatoms may also grow in profusion on other plants and animals. The littoral genus Licmophora frequently occurs on pelagic copepods, and the massed growth of Cocconeis ceticola flourishing on the skin of whales that have spent considerable time in the cold antarctic waters has, by its yellow color, given rise to the name “sulphur-bottom” for the blue whale.

Dinoflagellata. These are frequently spoken of collectively as the dinoflagellates (fig. 74). Space will not permit the amount of discussion that this diverse group of organisms requires for adequate treatment (see Kofoid and Swezy, 1921, Kofoid and Skogsberg, 1928, Fritsch, 1935). It is a group concerning which it is not easy to make generalizations without the danger of introducing errors. The members are of great importance in the economy of the sea. A large number are holophytic and rank second to the diatoms as producers in the marine plankton. They are therefore best studied with the phytoplankton. Others are holozoic or animal-like in nutritional requirements, ingesting particulate food and possessing other characteristics that place them clearly with the animals. Some are saprophytic, living upon dead organic matter. All are important as food to filter- and detritus-feeding animals.

Typically, the dinoflagellates are unicellular, some being armored with plates of cellulose, others unarmored or naked. All possess two flagella for locomotion, an important feature in the holophytic forms, for

they can thus, perhaps, adjust themselves in some degree to water strata most favorable with respect to light and dissolved plant nutrients. Many dinoflagellates are luminescent, being responsible for much of the brilliant luminescent display so characteristic of the sea.


Dinoflagellates and other phytoplankton organisms. a, Ceratium tripos; b, Dinophysis; c, Ornithocercus; d,e, Triposolenia, front and side views; f, Peridinium; g, Amphisolenia; h, Goniaulax; i, Ceratium fusus.

Methods of Reproduction. Among the dinoflagellates, reproduction is accomplished mainly by processes of cell division, which in some instances result in a chain of individuals clinging loosely together. Temporary structural variations may normally occur in individual cells at opposite ends of the chain. The progressive size reduction that is

characteristic of one daughter cell resulting from simple division in the diatoms does not occur after these divisions.

Dinoflagellates are found in all seas, but the greatest development of species is met with in the warmer waters, where a number of very bizarre forms are to be found. Owing to the destructibility of their cellulose plates by bacteria and other agencies, they are not preserved in bottom deposits. Important genera are Ceratium, Peridinium, Dinophysis, Gonyaulax.

Phaeocystis. Phaeocystis is a brown, flagellated plant, neritic in habit, that forms colonies in gelatinous, lobed globules visible to the naked eye. The large numbers produced may at times render the surface water quite brown and become a serious cause of clogging in silk plankton nets. Reproduction is accomplished by formation of flagellated spores that escape from the colonies.

Coccolithophoridae. Among the smallest (5 to 20 microns) of autotropic organisms of the sea are the biflagellated (some marine forms are not flagellated) forms of this group (fig. 74). Usually they are not caught by the ordinary net, through the meshes of which they readily escape, and when caught special care must be taken that their calcareous protecting armor is not dissolved by the preservative, leaving only an indefinable mass. The soft parts are shielded by tiny calcified circular plates or shields of various design and projections called coccoliths, or rhabdoliths. These shields had been found in enormous numbers in marine bottom deposits before the organisms of which they are a part were discovered by the Challenger and identified from their living habitat in the plankton, where they were found entangled in protoplasmic strands of pelagic protozoa or in the stomachs of salps and pteropods. Typically, the coccolithophoridae belong to the open sea, but they may occasionally reproduce in large numbers in coastal waters; at one time, according to Gran (1912) numbers of 5 to 6 million per liter gave the waters of Oslo Fjord a milky appearance. Some also occur in fresh water.

Though minute in size, they are of great importance as food to filter-feeding organisms, and also as contributors to calcareous bottom sediments. They occur in geological formations dating from the Cambrian period. Common genera among these organisms are Coccolithus, Pontasphaera, and Rhabdosphaera.

Halosphaera. Halosphaera is a unicellular, microscopic plant of the order Heterococcales (fig. 74). Earlier authors have included it with the green algae. It occurs at times in vast numbers in the plankton, floating mostly near the surface. Halosphaera virides occurs over the whole Atlantic and is abundant both in the warmer waters of the Gulf Stream system and in high southerly latitudes, where the Discovery investigations in the Antarctic found it second in importance to the diatoms. Meringosphaera of this order also occurs in marine plankton.


According to Gran, Halosphaera is practically the only open-sea form in which the predominantly green color of land plants is to be found. Notwithstanding the vast numbers that are often found, it does not reproduce by the quick method of simple binary fission, as in diatoms, but, after having grown for some time to its maximum size, the cell contents are transformed into a large number of zoospores. These swimming spores escape and through some unknown method are transformed back into tiny globular forms that gradually increase to normal size by successively shedding their weakly silicified investing membranes. Resting spores may also be produced.

Silicoflagellates. These flagellate organisms (fig. 74) deserve mention only briefly, since they do not usually occur in sufficiently large numbers to enter materially into the economy of the sea. However, they are such persistent members of plankton communities from nearly all colder seas that their starlike, open, siliceous shells attract a good deal of interest. Many occur in bottom sediments, and their development is shown in fossil marine deposits. That they contribute at least in a small way to the food of animals is shown by their frequent occurrence in food vacuoles of tintinnids.

The Higher Plants in the Sea

The two intermediate phyla of the plant kingdom—namely, the mosses (Bryophyta) and the ferns (Pteridophyta)—are wanting in the sea. However, the highest of plants, the Spermatophyta, are represented by about thirty species of Angiosperms, or flowering plants. These belong to three genera of the Hydrocharitaceae and six genera of the Potamogetonaceae (Arber, 1920). They have not originated in the sea, but have invaded and colonized it by way of fresh water. Their closest affinities are with widespread fresh-water angiosperms belonging to the same families.


The eel grass Zostera, to show leaves, rhizome, and true roots.

Of outstanding importance among the marine angiosperms is the eel grass, Zostera (fig. 75). Botanically, the plant is not a grass despite the long, slender, and flexible grasslike leaves, which are thus adapted to withstand the force of moving water. Unlike the benthic algae, Zostera and its relatives possess true roots that are attached to an underground stem, or rhizome, forming an anchor in the soft substratum. There are fertile and sterile plants, and, since the plants grow submerged, mostly in depths of 4 or 5 m but also to a depth of 14 m (Petersen, 1918), the flowers are pollinated under water through the agency of currents.

The threadlike pollen grains have a density near that of water, and are therefore readily carried about. The plants are also perennial, the rhizomes growing longer and giving rise to new sets of leaves and roots. Zostera marina has a wide distribution, occurring on the coasts of Europe, North America, Asia Minor, and eastern Asia. It grows best in coastal areas that are protected from violent wave action. Phyllospadix, a related genus, is confined to the open, wave-washed shores of the west coast of North America. Other genera and species are found in many parts of the world. In some extensive shallow-water areas—for instance, in the fjords of Denmark—Zostera is considered to be the main source of detrital food for marine animals.


Many different species of distinct animal groups live intermingled in the same faunal area. Some may have identical habits and requirements, but for the most part the separate species or higher ranks have characteristic limitations, and each has its own function in conditioning the whole complex organic environment, thus influencing the type of species forming the association. This will become increasingly clear in considering the interrelations of the organisms (chapter XVIII).

For an adequate understanding of the intricacies of the fauna, it becomes necessary, therefore, to understand the part played by the separate species or groups of species, and, in order to circumscribe and to interpret the geographic and bathymetric distribution of species, their exact identity must be established. Systematic biology—that is, taxonomy—provides the tools for these purposes and is therefore an indispensable aid toward the desired goal. The field of study is so overwhelmingly large, however, that the many species comprising the primary groups must be investigated in special studies and the results of many different specialists must be integrated to provide a picture of the fauna as a whole. Much still remains to be done in this field, for many sections of the oceans have been only superficially investigated.

A full descriptive treatment of the animals of the sea would require several volumes, but it will suffice for our purpose to list or succinctly to review the primary divisions, including only a few of the secondary divisions that in the development of the study of marine biology and oceanography have assumed more or less importance and that are illustrative of the general field.

In the following synopsis, where the number of species is recorded for any group, the figures have been obtained mainly from Pratt (1935) and Hyman (1940). Illustrations are found mainly in chapter XVII.


Synopsis of the More Important Systematic Groups of Marine Animals



Protozoa are single-celled organisms microscopic or minute in size. The sea bottom harbors many creeping and attached protozoa of the ameboid or ciliate types, but we shall be concerned mainly with the pelagic forms inhabiting the plankton.


Order Dinoflagellata. In the broadest sense, this group contains both animals and plants, it being a borderline group.

Foremost among the protozoa in the economy of the sea are the dinoflagellates, chiefly because of the capacity of many types to carry on photosynthesis. These holophytic members are considered more fully in the discussion on plants, and for oceanographic studies are properly included in the phytoplankton. It will suffice to mention here only Noctiluca (fig. 225g) as an important representative of the holozoic members, none of which have chromatophores. The soft spherical body of Noctiluca is pale pink in color and bears a conspicuous flexible tentacle. The maximum size is only about 1.5 mm, but, when reproducing in profusion by simple cell division, the countless numbers produced may, by their accumulation, impart a pinkish-red color to considerable areas of surface coastal water, and the masses may be blown into conspicuous windrows or patches resembling “tomato soup.” Noctiluca are voracious feeders, engulfing particulate food such as diatoms and other small organisms. This form is also important as a contributor to the luminescence of the sea.


Order Foraminifera. The oceanographic interest of this order (and also, to some extent, of the following order) lies in the skeletal structures produced by its members. In the foraminifera the shells are variously formed, with one or more chambers arranged in a straight line or in a spiral (fig. 225a). Some are provided with many pores for the projection of protoplasmic pseudopodia used in capturing food. The shells are constructed typically of calcium carbonate, but silica and chitin are also used, and in some

benthic arenaceous forms they consist of an agglomeration of foreign materials cemented together. The greatest numbers of individuals are planktonic in life, but upon sinking to the bottom the shells form an important constituent of globigerina ooze, named for the abundant pelagic genus Globigerina (p. 816). Fossil foraminiferal shells are much used in the study of geological strata, being indices of past geological conditions, and are useful in the field of oil geology. A few foraminifera occur in fresh water, but the vast majority occur in the sea, either living on the bottom even at depths of 6000 m or floating freely in the water, preferring the warmer seas. There are over 1200 species, of which about 26 are pelagic. A recent catalogue of foraminifera (Ellis and Messina, 1940) includes some 18,000 living and extinct species.

Order Radiolaria. These are planktonic organisms whose skeletons are composed mainly of silica, but the Acantharia contain acanthin (strontium sulphate), and all types possess an inner capsule of chitin. The siliceous skeletons are formed in the most intricate and widely divergent patterns in the different species and are the most beautiful of all objects found in the sea (fig. 225e,f). Upon sinking and mingling with the bottom sediments, the skeletons become the type constituents of the siliceous radiolarian oozes found most abundantly covering the ocean floor in the deep tropical waters of the Pacific Ocean (fig. 253). There are about 4400 species, all marine.


Suborder Tintinnoinea. These protozoans, commonly called tintinnids, are mostly of extremely small size, varying from 20 μ for Tintinnopsis nana to 640 μ for Cymatocylis robusta. Swimming is accomplished by the beating of a whorl of hairlike cilia at the anterior end. Their loricae, or shells, range in shape from tubular to urn-shaped structures that are secreted in a stereotyped fashion by the animal and may or may not include agglomerated foreign material such as bits of sand, diatom shells, and coccoliths (fig. 225c,d). The tintinnids at times are found in vast numbers, especially in coastal water, where they are important feeders on the smallest plankton, the nannoplankton. Their sensitivity to small changes in environmental conditions makes them fluctuate in numbers with seasonal or other changes. There are 692 known species,

mainly marine. (Kofoid and Campbell, 1929.) Examples: Favella, Tintinnopsis, Tintinnus.


The sponges are multicellular animals, though of simple and loose organization, either with spicules of silica or calcium carbonate imbedded in their bodies for support or with fibrous skeletons made of the horny substance spongin, as in the common commercial sponge. Sponges are all benthic and nearly all marine, only one family occurring in fresh water. In the sea they are to be found in all parts and at all depths, the siliceous forms living largely in the deep sea. Sponges grow attached to the substratum and obtain their food by propelling water through tiny pores in the body wall and filtering out the microorganisms and detritus that may be present. There are about 2500 species, mostly marine.


Coelenterata are tubelike primitive forms with a continuous body wall surrounding a simple digestive cavity with but one opening encircled by tentacles used in capturing food. The group shows a remarkable degree of polymorphism; that is, a single species may present a variety of forms reducible either to the sessile polyp or the swimming medusoid type.

Class Hydrozoa. To this class belong the hydroids commonly found growing in little tufts on rocks and sea weeds along the coast. From these branching polyps are budded the small jellyfish or medusae such as Obelia (fig. 79). The Siphonophora, an order of this class, are characteristic of the open sea and are represented by the beautiful blue Velella (“by-the-wind sailer”) (fig. 226b) and Physalia (the “Portuguese man-of-war”), neither of which possesses a sessile stage. They are planktonic colonial medusae, exhibiting the maximum development of polymorphism of all animals. There are about 2700 species of hydrozoa.

Class Scyphozoa. To this class belong the larger medusae with eight notches in the margin of the bell. Here are included the giant jellyfishes, some of which may become 2 m in diameter. A much-suppressed sessile polyp stage is present in the group. The 200 species are entirely marine. Examples: Aurelia, Cyanea.

Class Anthozoa. To this class belong the sea anemones, corals, and alcyonarians. There is no medusoid stage, and many of the polyps are colonial; some, especially the corals, are notable for their precipitation of calcareous skeletal structures, which, through long periods of accumulation, are important in the

building up of coral reefs and similar formations. All 6100 known species of anthozoa are marine.


Ctenophora are small globular or flattened forms of jellylike consistency and with eight meridional rows of fused cilia used in swimming. Some possess a pair of trailing tentacles used in the capturing of food. The abundant globular species are commonly known as “comb jellies” or “sea walnuts” (fig. 226a). There are 80 species, all marine. Numerically important genera are Pleurobrachia and Beroë.


Platyhelminthes are flatworms, a large number of which are found in the sea, either free-living or parasitic.

Class Turbellaria. Nearly all of this class are free-living on the bottom under stones and in crevices, where they move about by means of cilia covering the body.

Class Nemertinea. These are ribbonlike worms sometimes considered as a separate phylum. The benthic species live among rocks, algae, mussels, and so on, or burrow in the bottom, where they capture small organisms by means of a long eversible proboscis. Extraordinary size variations occur, some species being only 5 mm long, while one, Lineus longissimus, may become 25 m in length when extended, and therefore is the longest of the invertebrates; however, its threadlike form contains but little bulk. Fifty-two planktonic species of nemerteans are known, some living at great depths—for example, Pelagonemertes. (Coe, 1926). The planktonic forms are modified, some with caudal and horizontal fins for swimming (fig. 228c). There are about 550 species of nemerteans, of which nearly all are marine.


The thread or round worms occur largely as parasites, but some are found in the plankton, and very large numbers occur in decaying organic detritus on the bottom. There are about 1500 species, many of which are nonmarine.


Class Rotatoria (Rotifera). These are tiny benthic or planktonic organisms provided with rings of cilia for swimming and for gathering food. Vast numbers may occur in the neritic plankton during the warmer seasons. There are about 1200 species of rotifers, of which most are fresh-water inhabitants.



These colonial animals, known as “sea mats” or “moss animals,” form flexible tufts or thin incrustations over the surface of solid objects both in intertidal and deep waters. Below low tide, many species form rigid, erect, latticed or branched colonies. The individual minute animals have calcareous protective skeletons and possess a ring of ciliated tentacles for gathering microscopic food. There are over 3000 species, about 35 of which are nonmarine.


Brachiopoda are ancient sessile animals superficially resembling bivalve molluscs, but the hinged calcareous or horny shells are dorsoventrally situated instead of laterally, as in the molluscs, and the animals gather their food by means of delicate ciliated arms attached within the shell. They grow permanently attached to rocks and shells, usually in the littoral zone below low tide. A few live in burrows. All are marine and all are very abundant as fossils in the Paleozoic and Mesozoic rocks. About 120 living and 3500 fossil species are known.


Phoronidea are wormlike animals, living in membranous tubes in the sand and collecting food by means of ciliated tentacles. There are about 12 marine species.


Chaetognatha include numerous but small (maxima about 75 mm long) holoplanktonic wormlike animals known as “arrow worms” or “glass worms.” They are highly transparent and provided with eyespots, a caudal fin and one or two pairs of lateral fins, and with strong chitinous jaws and teeth for capture of prey. They occur from the surface to great depths and are distributed far to sea in all latitudes. All 30 known species are marine. Sagitta (fig. 228a) is the most abundant genus.


Annelida are true worms with elongated bodies composed of a series of similar segments.

Order Polychaeta. These are marine worms of great abundance provided with many setae and typically with a variety of well-defined head structures such as eyes, tentacles, chitinous jaws, ciliated cirri, and so forth, which are modified in keeping with their habits of life and mode of feeding. They have a wide distribution horizontally and bathymetrically.

For the most part they are benthic, either creeping or burrowing, as in Nereis, Glycera, and Arenicola, or sedentary in calcareous or fibrous tubes, as in Spirorbis and Sebella. Tomopteris is pelagic (fig. 228b). There are in all over 3500 species, nearly all marine.

Order Oligochaeta. These are earthworms, of which only a very few are marine, living near shore.

Class Echiuroidea. These are fleshy marine worms with only one or two pairs of setae. They are unsegmented or indistinctly segmented in the adult. They live in burrows in the mud and sand of the littoral zone. There are about 20 species.


Arthropoda include animals with a segmented, chitinous exoskeleton and with jointed appendages, variously modified for locomotion, feeding, and other activities.

Class Crustacea. Entomostraca. This group, formerly considered a subclass, is of convenience in designating a large assemblage of small, primitive crustacea belonging to several subclasses and orders distinguished from the higher crustacea, or Malacostraca.

Suborder Cladocera. Only a few occur in the sea. Examples: Podon, Evadne, sometimes important in neritic plankton. Very numerous in fresh water.

Order Ostracoda. This order includes more than 2000 species, mostly marine, living in the plankton and on the bottom (fig. 227b).

Order Cirripedia. These are the barnacles which as adults have calcareous shells and live sessilely in all benthic habitats, especially coastal. Some grow attached to drifting objects or upon whales and other animals, or they may form special floats for suspension. There are about 500 species, all marine.

Order Copepoda. Though small in size (about 0.3 mm to 8 mm in length), the copepods bulk large in the animal substance of the sea, for they are by far the most abundant of all crustaceans and usually constitute about 70 per cent of the zooplankton. There are over 6000 species of copedods, found mostly in the sea, where some 750 species are planktonic and extremely numerous. Many others are benthic or parasitic. The three main suborders of free-living forms are Calanoida (fig. 227c), Cyclopoida (fig. 229d), and Harpacticoida (fig. 229a). The first two are mainly pelagic,

the last benthic. Like other Entomostraca and some Malacostraca, they gather food by means of fine bristles on certain appendages (p. 887).

subclass malacostraca. These are the large crustacea, mostly benthic, many with strong claws and biting mouth appendages.

Order Mysidacea. There are about 300 species, mostly marine, living on or near the bottom.

Order Cumacea. About 400 species of this order are known; nearly all are marine, benthic.

Order Euphausiacea. These are commonly known as “krill,” and in some regions are very abundant in the plankton and near or on the bottom. Some attain a length of about 50 mm, and may at times be the major constituent of the zooplankton. There are 85 known species, all marine. Examples: Euphausia, Meganyctiphanes (fig. 227a).

Order Amphipoda. There are about 3000 species, nearly all marine, in various habitats.

Order Isopoda. Over 3000 species are known; they are mostly marine, living on the bottom and on vegetation or burrowing in wood. Examples: Limnoria, Munnopsis (figs. 77 and 221).

Order Stomatopoda. This order contains about 200 species, all marine, benthic, most common in shallow water of lower latitudes.

Order Decapoda. Decapoda include crabs, lobsters, shrimps. They are widely distributed in both the pelagic and benthic regions. Most of the over 8000 species are marine.

Class Arachnoida. This class is well represented in the sea by a number of marine mites, over 400 species of sea spiders or pycnogonids, and 5 species of Limulus, the king crab. All are benthic.

Class Insecta. Only one insect is submarine during its whole life; a few others live on the foreshore or skip over the surface in search of food. Example: Halobates.


The molluscs are noted particularly for their construction of an infinite variety of calcareous shells encasing the body and for the structural modifications that have taken place in the soft parts known as the foot and the mantle. These modifications are associated with the method of locomotion and capture of food.


Class Amphineura. The chitons are all flat, benthic animals creeping with the aid of a broad, flat foot. There are about 630 species, all marine.

Class Scaphopoda. Tusk shells live in the bottom mud from shallow water to depths of over 5000 m. All 200 known species are marine.

Class Gastropoda. In most types there is a spiral shell, and the foot is used in creeping. In this and the preceding classes a rasplike radula is a characteristic food-gathering organ. Some gastropods are holoplanktonic and may be without shells. These are the marine pteropods and heteropods (about 90 species of each) with the foot modified for swimming (fig. 228d,f). The latter are especially characteristic of the oceanic waters of the lower latitudes. There are about 49,000 species in the class, mostly marine.

Class Pelecypoda. The clams, oysters, and mussels have a hatchet-shaped foot which in many is used for digging. All are benthic, usually sessile or burrowing in mud, rock, or wood. The soft parts are enclosed within hinged shells and the food is conveyed to the mouth by means of ciliary action setting up water currents, sometimes through long siphons. There are about 11,000 species, of which about four fifths are marine.

Class Cephalopoda. In the squids, devilfish, and so forth, the foot is divided to form arms used in capture of prey. In keeping with their active, predacious habits, the eyes are usually well developed, but blind deep-sea forms occur. In Nautilus and related forms there is a well-developed shell. Cephalopods are either benthic or pelagic, some living at great depths. The giant squid, Architeuthis princeps, having a body girth of nearly 1 m and attaining a total length of about 16 m, is the largest of all invertebrates. There are about 400 species, all marine.


Echinodermata are animals with calcareous plates forming a more or less rigid skeleton, or with scattered plates and spicules embedded in the body wall. Many are provided with spines. All are marine, and all but a few sea cucumbers are benthic.

Class Holothuroidea. The sea cucumbers are mainly benthic. only members of the order Pelagiothurida being planktonic. There are over 650 species, some living in abyssal regions.

Class Asteroidea. The sea stars are among the most conspicuous of shore animals, but they live also at very great depths. About 1100 species are known.


Class Ophuroidea. There are more than 1600 species of brittle stars, with a wide horizontal and bathymetric distribution.

Class Echinoidea. There are about 600 species of sea urchins and sand dollars, a few of which live in deep water.

Class Crinoidea. About 800 species of sea lilies and sea feathers are known, with the center of distribution in the East Indian waters, but they also occur in many other waters. The former live mainly in the deep sea and are anchored by long stalks. The latter occur mainly at shallower depths and are without stalks. The class is a vanishing remnant of a formerly abundant group that has left more than 2000 fossil species.


Chordata are animals which in some stage of their life have gill slits and a skeletal axis known as a notochord.

Subphylum Tunicata. These are primitive chordates; of about 700 species, all are marine.

Class Larvacea (Appendicularia). These are small planktonic forms, sometimes abundant. Examples: Oikopleura (fig. 228e), Fritillaria.

Class Ascidiacea. These are sessile ascidians such as Ciona and Culeolus.

Class Thaliacea. This class is made up of pelagic tunicates that float singly or in chains; they may be very abundant at the surface in the warmer waters. Examples: Salpa, Doliolum.

Other protochordates are the wormlike Enteropneusta and the fishlike Cephalochorda, both of which are found burrowing in mud and sand.


Subphylum Vertebrata. This group includes animals with vertebrae. All but the classes Aves and Mammalia are cold-blooded.

Class Cyclostomata. The hagfishes and lampreys are fishlike forms but without paired fins. They have a circular sucking mouth without jaws. The former are all marine, while the latter live both in the sea and in fresh water.

Class Elasmobranchii. These primitive fishes—the sharks, rays, and chimaeras with a cartilaginous endoskeleton—have paired fins and a lower jaw. In this group are many large forms such as the giant manta and the whale shark, the largest of all fishes, which becomes about 16 m long. Nearly all are marine.

Class Pisces. This class includes the true fishes, with a bony endoskeleton, paired fins, and an operculum covering the gills. They are characteristically streamlined for great swimming

speed, but a considerable variety of structural modifications occurs. Like the above class, they are mostly carnivorous and highly rapacious. Most fishes are marine, and some are benthic, but the majority are pelagic, living in both shallow and abyssal depths.

Class Reptilia. This class is represented in the sea by snakes and turtles. They breathe air and are therefore inhabitants of surface waters. The turtles frequent the shore to deposit their eggs on sandy beaches; the snakes bring forth living young and are therefore less dependent upon the shore. The sea snakes are found in the Indo-West Pacific and in tropical waters of America. They grow to a length of from 1 to 2 m or more and some are very poisonous. The sea turtles occur in tropical and subtropical seas. They have paddlelike limbs for swimming, and some grow to great size. The leathery turtle, for example, which is the largest of the class, may attain a weight of 1000 pounds.

Class Aves. A great number of birds are dependent upon the sea for food. Some of these frequent the land only for nesting and rearing of young. Typical examples are the albatrosses, petrels, cormorants, and auks.

Class Mammalia. These are warm-blooded, air-breathing animals with hair and mammary glands.

Order Carnivora. The marine members of this order are the sea otters and, to a lesser degree, the polar bears. The sea otters occur only in small numbers and only along the west coast of North America, where they were formerly hunted commercially to the very verge of extinction. Recently, under rigid protection, they have recuperated to an encouraging degree. The polar bears are confined to the Arctic region, usually on or near floating ice

Order Pinnipedia. Pinnipedia include seals and walruses, nearly all marine. The limbs are finlike, in adaptation to the aquatic existence. There are three families: (1) Otariidae include the eared seals, sea lions, and fur seals. Small external ears are present and the hind limbs can be rotated forward. (2) Phocidae are the hair seals without external ears and with hind limbs incapable of rotation forward. (3) Odobenidae include the walruses, with greatly elongated canine teeth in the upper jaw. They are confined to the Arctic.

Order Sirenia. Sirenia are heavy-bodied mammals with a flat tail and with forelimbs modified as paddles. Hind limbs are wanting. They live near shores in warm waters,

where they browse upon vegetation. They are not numerous. Examples: sea cows, manatees, and dugongs.

Order Cetacea. This order includes whales and dolphins, highly modified for aquatic life by a streamlined body and finlike forelimbs and tail. The hind limbs are wanting.

Suborder Mysticeti. These are the baleen, or whalebone, whales, with a series of long plates of baleen suspended in the mouth (fig. 76a). The frayed ends of these are used in screening out plankton food. Examples: fin whale, humpbacked whale, and blue whale. The last named is the largest of all animals, growing to a maximum length of about 34 m and weighing 294,000 pounds.


a, the blue whale—a whalebone whale; b, the sperm whale—a toothed whale.

Suborder Odontoceti. Odontoceti are the toothed whales. This group includes (1) sperm whales with teeth only in the lower jaw (fig. 76b) and (2) the numerous dolphins and porpoises with teeth in both jaws.

Reproduction and Life Cycles in Marine Animals

In any comprehensive study of oceanography wherein biological activities are implicated or in the study of any population or individual species in relation to environmental factors, it is necessary to take into consideration the nature of the life cycles of the organisms involved. Only thus can the biological activities and the methods whereby the race is maintained through countless numbers of generations be fully understood. The utility of life-cycle studies in such practical fields as economic entomology, parasitology, and fisheries has been abundantly demonstrated. Through a knowledge of the methods of reproduction and through recognition of the various developmental stages of marine animals the investigator has at hand valuable means of aiding the interpretation

of the fluctuations that occur in a given population, of understanding the vertical or horizontal migrations, and of tracing the methods and routes of distribution, because all of these phenomena are closely bound to phases in the life cycle of the organisms.

It should be pointed out here that, in dealing with the propagation of any individual species in relation to its distribution, we must distinguish broadly between (1) reproductive distribution and (2) sterile distribution. Reproductive distribution is associated with areas where environmental conditions are favorable to maturation, spawning, and larval development. Such areas may be called areas of reproduction, or nursery areas. Sterile distribution is associated with areas in which the submature or adult individuals may live and some spawning may take place, but in which the eggs fail to hatch or the larvae do not survive, so that the area must be restocked periodically by invasion of postlarval stages produced elsewhere.

In a study of the life cycles of marine animals, one is impressed particularly with three facts: (1) the preponderance of animals which, though sessile, creeping, or burrowing in the adult stage, possess a free-swimming period during the early stages of life; (2) the enormous numbers of young that are produced by both pelagic and benthic animals; and (3) the fundamental similarity of the larvae of different invertebrate groups. We shall be concerned only with the first two.

In a superficial survey of populations, it is mainly the larger, more conspicuous adult animals that are seen, yet, from the standpoint of numbers, vastly more starfish, barnacles, clams, crabs, fish, and so on, are represented in the microscopic, feebly swimming larval stages than in the adult stages. Most of these larvae do not survive to assume the adult habit, but, instead, serve as nourishment for other organisms, swimming or sessile, or are in some manner destroyed through action of the physical or chemical environment.

Types of Reproduction. In reproduction, animals are either oviparous or viviparous. The oviparous forms deposit eggs that develop outside the mother's body, while in the viviparous forms the young are nourished by the mother and are born alive in a postembryonic state. An intermediate condition exists in the ovoviviparous forms, where the eggs are incubated and hatched within the body, as in certain sharks, perch, and blennies. The term larviparous is sometimes used to indicate that larval stages are born. An embryo derives its nourishment from the yolk of the egg or directly from the mother, whereas typically a larva is morphologically adapted with mouth and digestive tract for the purpose of seeking its own nourishment. Later it will be seen how important this fact is in the life of many marine animals.

By far the greater number of animals of the sea are oviparous, and it is among these that the extraordinarily large numbers of eggs are

produced. The number of eggs produced by the female of any species is associated, however, with the degree of parental care or other protection accorded the eggs and larvae following fertilization. The greater the care, the fewer the eggs produced. Most spawning consists of casting the eggs freely into the water, where they are fortuitously fertilized by spermatozoa that have also been extruded into the water. In these instances, enormous numbers of eggs are shed each breeding season. The following examples of the numbers of eggs produced by single individuals will illustrate the extraordinary fecundity that is attained:

American oyster 115,000,000 Pacific halibut 3,500,000
Sea hare (Tethys) 478,000,000 Cod 4,400,000
Teredo navalis, more than 2,000,000 Sunfish (Mola) 300,000,000

Parental care of eggs and larvae.

It has long been recognized, however, that the exceedingly great number of eggs produced by some species is not directly correlated with the number of adults that are found. The large numbers of eggs and larvae produced are, instead, a measure of the tremendous toll paid by these species in order to assure survival of enough individuals to carry on the race.

In the marine population as a whole, very little parental protection is given to the offspring in the larval stages, and frequently even the eggs are given no care, yet hundreds of examples can be cited wherein varying degrees of protection are afforded the embryonic stages and sometimes the larvae as well. Many of the larger crustacea retain the developing eggs attached by secretions to hairlike structures on the abdominal appendages. Some annelids produce viscid secretions for attaching the eggs to setae or to the body wall, while others retain the young up to a well-developed larval stage in special brood pouches, as in Spirorbis (fig. 77). Many other invertebrates provide brood pouches—for example,

the isopod wood gribble, Limnoria, in which chitinous flaps cover the eggs and young. In many copepods the developing eggs are retained in membranous sacs (fig. 77). Among the vertebrates may be mentioned the male pipe fish, which carries the developing young in a special groove on the ventral side of his body. Other animals guard their eggs by hovering over them, as in the slipper shell, Crepidula, or the little six-rayed starfish, Leptasterias hexactis, which continues protection even beyond the larval stage. Many other examples of parental care are discussed by Wilson (1935).


Some types of egg cases for protection of eggs and larvae.

It is clear that in these instances of parental protection the need for large numbers of eggs is somewhat diminished. Nevertheless, when a relatively long, helpless, pelagic stage follows the protected period of incubation, many larvae must still be produced. Thus, for instance, the blue crab, which, though protecting the eggs till the young are hatched, has pelagic larvae and is said to carry over two million eggs (Truitt, 1939). In contrast, Limnoria produces a maximum of only about-twenty-five eggs, but retains these within a pouch until the young are able to burrow into the wood where they were born. Thus they escape the hazardous pelagic life of larvae. In this animal the hazards of a swimming existence are met not by the very young but by submature specimens which by short migrations attempt to establish themselves in less crowded situations prior to breeding (Johnson, 1935). The fact that some animals produce more eggs than others and at the same time offer more parental care must indicate that factors operate to destroy more developing young in one than in the other.

Finally should be mentioned the common method of depositing eggs in masses or protective capsules of various types (fig. 78), thus diminishing loss from excessive dispersal and other hazards of a floating existence during the period of incubation. The capsules are also sometimes

watched over in the manner of the octopus, which keeps her eggs free from enemies. Additional examples are given in the discussion of life cycles.

Types of Development. There are two main types of development: (1) direct and (2) indirect. The direct development in oviparous species is associated with eggs of considerable yolk content, such as are found in the fishes, cephalopods, some nemerteans, crustaceans, and others. The newly hatched young are similar to the parent except for size.

Direct development is common among the deep-sea benthic animals, and this habit appears to be an advantageous adaptation. The slowly moving currents of great depths are of less importance in the dispersal of larvae than are the stronger currents of shallow water. The micro-planktonic life so characteristic of surface layers and from which the pelagic larvae of littoral animals directly derive their food has no counter-part in the deep, and hence it is imperative that the young produced be able to feed directly upon the bottom detritus. The possibility of deep-sea larvae swimming from great depths to the surface, where food is plentiful, and later returning to the bottom appears to be impracticable in nature. The young of some mid-depth pelagic forms—for example, Cyclothone among the fishes and Acanthephyra among the prawns—do however live nearer the surface, where food is more plentiful, than where the adults are commonly to be found (Hjort, 1912).

Benthic animals of Arctic and Antarctic regions also commonly possess no pelagic larval stages. Hjort (1912) and Murray (1913) consider this a probable explanation of the great local concentrations of certain boreal and arctic benthic animals, because the direct development results in the young remaining in the area in which they are born. Brief pelagic larval stages following protection during incubation and absence of dispersing currents lead also to local adult concentrations.

The indirect development is associated with a type of egg with little yolk (that is, alecithal), and hence a self-sustaining larva must develop quickly or the organism dies. This type of development is characteristic of marine invertebrates, which usually cast their eggs free in the water or carry them through the incubative period in special brood pouches. Larval stages appear before the full character of the species to which they belong becomes established. Many of these—for example, the pluteus larvae of the Echinoidea and the Ophuroidea—when first discovered were described as distinct kinds of animal, only to be found later to be the young of already well-known species. The locomotor organs of most of the larvae are cilia (see below for exceptions) which by their rhythmic beating propel the animal slowly through the water at a rate just sufficient to keep them in suspension. The great similarity of structure exhibited by the larvae of some groups suggests a common origin for the groups which as adults are structurally very dissimilar.


Typical Life Cycles. The life cycles of many species have not been investigated, but the principal features in the life history of the major groups have been established. We shall here review only the groups of most immediate interest in general oceanographic studies.

In the protozoa, reproduction is mainly by binary fission, whereby the animals divide to form two separate animals, these in turn dividing after growth. Under favorable conditions, this method makes possible a production of great masses of individuals, as is often witnessed in such forms as Noctiluca. Gametes are also formed in this animal as a result of multiple fission. These unite in pairs, but their further development is unknown. In foraminifera, and possibly also in radiolaria, there is a cyclical alternation of generations in which sexual and asexual phases alternate and give rise to morphologically different individuals (Myers, 1936).

In the tintinnids, in which transverse binary fission occurs, the anterior daughter escapes from the lorica, while the posterior daughter retains the old lorica (Kofoid, 1930).

The sponges reproduce asexually by budding or fragmentation, and sexually by union of gametes, the latter resulting in a free-swimming, flagellated larva, the amphiblastula, which, after a period of swimming, settles to the bottom and grows to form the adult sponge. Asexually produced units, known as gemmules, possessing a heavy protective covering are produced by some sponges as a means of survival during adverse periods. Reproduction by formation of gemmules occurs principally among the fresh-water sponges, but some marine forms also produce gemmules.

In the coelenterates, both sexual and asexual reproduction are important features in the life cycle. The union of germ cells results in a free-swimming, ciliated, planula larva about 1 mm long (fig. 80c). The planulae, though lacking the mouth and enteron of typical larvae, may live for a sufficiently long period on yolk food to bring about dispersal of sessile coelenterates such as the corals and anemones. Vaughan (1919) found the pelagic period of corals to be from one day to two or three weeks. Upon settling to a hard bottom the planulae of corals and other Anthozoa develop a mouth and tentacles for feeding, and later the reproductive organs are formed. In some there is also active asexual reproduction by fission and budding. Large coral colonies are thus initiated from a single individual. It is the skeletons of these asexually produced individuals which form the large coral heads, some of which are 3 m or more in diameter and contain many thousands of individual polyps. The length of time required for formation of such colonies has been investigated by Vaughan (1919), who found that a coral colony (Porites asteroides) 50 mm in diameter may be formed in four years.


The remarkable life histories of many jellyfish of the class Hydrozoa offer the examples of alternation of generations that are used in all zoological texts. The jellyfish, or medusa, stage (fig. 79) of such forms as Obelia is either male or female, and the eggs are cast free in the water, where they are fertilized and develop into planula larvae. The planula soon settles on the bottom to form the sessile polyp, or hydroid, stage. From special structures on the polyp, asexually produced buds become separated as swimming medusae, thus completing the cycle. The alternation of sessile and pelagic generations is a factor of great significance in the distribution not only of the hydrozoa such as Obelia, but also of other forms—for instance, the large scyphozoan Aurelia, which, though varying in details of life history, possesses stages similar to those in Obelia. The sessile generation is the chief link instrumental in restricting all stages of such animals to the neritic waters, generally to the proximity of shores and shoals with a suitable substratum of rocks, shells, or larger plants for attachment of the planula larvae. Bigelow (1938) found that in the open sea off Bermuda only about 3 per cent of the medusae caught at a distance of 10 miles from shore were of the type with a fixed stage in their life history. The degree of dispersal is, of course, dependent upon the speed and direction of the water currents prevailing. Some swimming jellyfish—for example, Aglantha digitalis and other members of the order Trachylina—are not dependent upon a sessile stage because daughter medusae develop directly from the pelagic stages.


The life cycle of a typical hydrozoan jellyfish, Obelia.

Open-sea colonial coelenterates—for example, Velella or Physalia—are representatives of the “blue-sea fauna.” Their life cycle is adapted to offshore life by elimination of the sessile stage. The planula larva gives rise to a medusiform stage from which the complicated colony arises.

The Ctenophores are all hermaphroditic, and the eggs are usually shed into the water, where, upon fertilization, they grow by direct development into free-swimming larvae. Gastrodes, a parasite in Salpa, produces a typical planula larva.

The great importance of annelids, especially in the littoral benthic fauna, warrants their inclusion in this brief study of life histories. At certain seasons the voracious swimming larvae of annelids are a major

constituent of the temporary plankton and a factor to be considered in the control of phytoplankton production in neritic waters at these times (p. 772). Most marine annelids are unisexual, in contrast to the hermaphroditic terrestrial forms. Both sexual and asexual reproduction occur, but, when great numbers of pelagic larvae are found in the plankton, they result from the shedding of many eggs free in the water, where they are fertilized and develop into ciliated, swimming, trochophore larvae (fig. 80b), which are soon transformed into miniature worms with three or more segments before they desert the plankton (fig. 224a) There are many modifications of the trochophore, its fundamental structure being reflected in the larvae of a number of animals, especially the molluscs, the nemerteans, and other flatworms.


Some characteristic marine larvae. a, nauplius larva of the copepod Labidocera; b, trochophore larva of the annelid Nereis agassizi; c, planula larva of coelenterates; d, zoea larva of the crab Pachygrapsus; e, veliger larva of the clam; f, bipinnaria larva of starfish; g, the cod larva with yolk sac.

The larvae of many benthic annelids enter the plankton only after they have completed their early stages under some means of special protection. For example, in the little tube worm, Spirorbis, and related forms, they develop in a special brood pouch beneath the operculum, while in some Polynoe, or scale worms, they are sheltered by the dorsal, flaplike elytra, and in yet other instances the eggs are deposited in attached or demersal gelatinoid masses (fig. 78), where the developing embryos and larvae enjoy some degree of protection. In Spirorbis the trochophore stage is passed in the brood pouch and the older larvae may assume the sessile habit after only twenty-four to thirty-six hours in the plankton. If a suitable substratum is not available, the pelagic stage is somewhat prolonged. Nereis agassizi spawns the eggs free

in the water, and the pelagic stage may be of twelve to sixteen or more days' duration. In N. vexillosa the eggs are spawned in masses (fig. 78).

The chaetognaths, or arrow worms, are hermaphroditic but not self-fertilizing. The eggs, which are fertilized internally, are shed into the water, where they develop directly into free-swimming larvae not unlike the adults.

Most crustaceans, in which we are particularly interested because of their prominence in some phases of oceanographic studies, pass through several distinct pelagic larval stages. The common initial crustacean larva is the nauplius (fig. 80a), bearing three pairs of appendages used for both swimming and feeding.

In copepods the sexes are separate, in some species the ratio of adult males to females being strikingly unequal at all times, while in other species the inequalities are seasonal, the males being more abundant at the onset of breeding but later diminishing in numbers more rapidly than the females after the breeding season (Damas, 1905, Farran, 1927, Campbell, 1934). The most important of the planktonic copepods spawn the fertilized eggs free in the water, yet many littoral species and some important pelagic species—for instance, Oithona, Paraeuchaeta, and others—carry the eggs in brood sacs through the period of incubation (fig. 77). In both cases the eggs hatch to typical self-sustaining nauplii. Paraeuchaeta is somewhat of an exception, for it develops from a heavily yolked egg and does not feed in the naupliar stage (Nicholls, 1934). In copepods there are normally six successive naupliar stages separated by definite moulting of the chitinous skin. The hard exoskeleton of crustaceans does not grow, and must therefore be shed or moulted periodically as the animal becomes too large for the encasement. In many crustaceans the number of moults may be variable, but in copepods there are a fixed number of stages, each separated by a moult. At the termination of the sixth naupliar stage a complete metamorphosis occurs from which emerges Stage I of six successive copepodid (copepodite) stages. Copepodid Stage VI is the adult, and, during spring reproduction in waters of the latitude of the British Isles, maturity may be reached in a period of about twenty-eight days in Calanus finmarchicus, but it is much delayed in the autumn-winter generation or in populations of more northern waters.

In Calanus finmarchicus, by far the most thoroughly investigated of all pelagic copepods, it has long been known (Gran, 1902, and others) that the animals spend the winter months in the deeper water layers. The breeding of this species occurs in spring and summer in boreal waters, and there are two or more successive generations, each of which, apparently, may bear more than one brood. The generation arising from the first spring spawning appears to mature quickly, spawn, and die. The last generation produced in autumn is a relatively long-lived

one, because it is this generation which carries the stock over the winter period. The winter stock is found in relatively deep water and is of uniform composition, consisting of copepodids in Stage IV or V. The stock is much reduced during the winter, but with the return of spring the animals pass into Stage VI (the adult stage) and spawn in surface waters, to produce the first generation of the season. Figure 81, from Nicholls (1933), shows changes in the percentage composition of the population, indicating three main periods of spawning in the Clyde Sea area.

The life cycle of Calanus finmarchicus appears to be quite characteristic of other members of that important genus and perhaps of other related genera as well, but very few pelagic copepods have been adequately investigated, and considerable variation can be expected. We shall not enter into the remarkable life histories of the parasitic copepods, but an example of a typically free-living littoral form—that is, Tisbe furcata—is instructive for comparison with Calanus. Tisbe furcata carries the eggs through the incubative period in brood sacs. There are six naupliar and six copepodid stages, as in Calanus but the time required for development from egg to mating maturity may be as little as ten days, and filled egg sacs are carried by females of the new generation in fourteen days after hatching. During maximum production, one female may produce at least seven or eight broods at about five- to eight-day intervals.


Successive generations of Calanus finmarchicus (from Nicholls).

In the euphasiids, another group of outstanding importance in the economy of the sea, the method of reproduction is not unlike that occurring in many copepods, in that usually the eggs are shed in the water, but the succession of generations is not rapid and the life span is of greater length. For Euphausia superba, in Antarctic waters, the time required to reach sexual maturity is estimated by Ruud (1932) to be two years. Some investigations indicate that normally the animals live in the immediate vicinity of the bottom and that during spawning they congregate in swarms and ascend to deposit the eggs in surface-water layers. Here, while slowly sinking, the eggs hatch to typical naupliar larvae, which are followed by successive stages of distinctive larvae, the older of which may return for a time to the surface layers. In

Thysanoessa inermis the stages are, in order, two nauplius, one metanauplius, three calyptopis, fourteen furcilia, twelve cyrtopia, and the adult (Lebour, 1926). Some species—for example, Nyctiphanes couchii—carry the eggs in a brood pouch through the incubative and naupliar periods.

Among most other crustaceans the eggs are carried through incubation attached to appendages or in brood pouches of various types. In heavily yolked eggs, as in the common crab and related forms, the nauplius stage is passed within the egg, and the developmental stage emerging from the egg is known as a zoea larva (fig. 80d), of which there may be several separate stages. The weakly swimming zoea may drift in the plankton for several weeks before changing to the megalopa stage and settling on the bottom. The lobster produces a special type of pelagic larva. the phyllosoma, which, with its leaflike body, is especially adapted to float in the plankton (fig. 229g). In barnacles, the larvae escaping from the mantle cavity within the shell of the sessile adult are typical nauplii which, after living a pelagic existence for a few weeks, are transformed into what are called cypris larvae (fig. 2241). The cypris larvae soon settle to the bottom and, upon attachment to a solid surface, metamorphose to assume the adult state.

In most bivalve (pelecypod) molluscs the sexes are separate, although some are hermaphroditic, and others—for example, certain species of oysters—are unisexual but change their sex alternately from one to the other. This phenomenon is also observed in other marine animals (Coe, 1940). Fertilization of the eggs frequently takes place after the eggs have been shed into the water, but commonly the eggs are retained within the brood pouch formed by the gills, in which case the spermatozoa are taken in through the inhalant siphon with the stream of water that is kept flowing over the gills. A modified trochophore larva is first formed, and from this a later larva, the veliger, results (fig. 80e). After a period of swimming the veliger settles to the bottom.

Gastropods are often hermaphroditic. Fertilization among them is commonly internal, and their eggs are frequently deposited in gelatinous or membranous cases attached to rocks or sea weed. Tiny floating cases containing several eggs are sometimes formed, as in Littorina (fig. 78). The trochophore and part of the veliger stage are passed within the egg case.

Among the echinoderms the sexes are separate and the eggs are usually spawned into the water, where fertilization occurs. Some (Asterina) lay demersal eggs which, because of their viscid nature, adhere to rocks and other objects. In other echinoderms, especially deep-sea and polar species, the eggs are fertilized and retained in brood pouches, where they undergo early development. Development is indirect in all cases, but the forms with most heavily yolked eggs do not have pelagic

larvae. Most echinoderms, however, do possess pelagic larval stages that may be of several weeks' duration.

The metamorphosis resulting in the adult state is as complete as that experienced by the butterfly, and it is not surprising that, before the parentage of the pelagic larvae was known, they were considered as distinct animals unrelated to the adult (figs. 80f and 224f and j). The characteristic larvae are bipinnaria (sea stars), echinopluteus (sea urchins), ophiopluteus (brittle stars), auricularia (sea cucumbers). The echinoderm larvae are of only moderate interest in the economy of the sea, but great biological interest is attached to the probable significance of some in showing a relationship to the most primitive chordates.

Many fishes—for example the cod, mackerel, halibut, and sardine—shed their eggs into the water, where fertilization takes place and the developing larvae are nourished by the yolk of the floating eggs (fig. 80g). The herring deposits viscid eggs which, upon sinking to the bottom in shallow water, become attached to solid objects. The gobies, blennies, sculpins, and others attach their eggs to solid objects or lay them on the bottom, where the male may stand guard over them until hatched. The grunion buries its eggs in the sand of wave-washed beaches during periods of high spring tide. Here the eggs remain for a period of about two weeks, when the next series of high tides washes them out and stimulates their hatching (Thompson, 1919, Clark, 1925). In sharks and rays, fertilization is internal, and either the young are born alive or the nonbuoyant eggs are deposited in leathery cases known as “mermaids, purses” (fig. 78, p. 317).

The eggs of fishes fall roughly into two groups, depending upon buoyancy: (1) pelagic and (2) demersal. The demersal eggs sink to the bottom or are deposited there; pelagic eggs float freely in the water and hence greater numbers are produced to overcome the losses inherent with this group. Many fisheries investigations are concerned with the occurrence and dispersal of pelagic eggs and the resulting larvae, for from such studies much information is gleaned regarding the spawning habits and areas of many commercially important fishes (p. 861). In general, the development of fishes can be considered as being direct, there being no general metamorphosis of form to the adult morphology. There is frequently, nonetheless, a marked degree of dissimilarity between larvae and adult, and in some there is a distinct metamorphosis. The Leptocephalus larva of the eel, for instance, was once considered a separate species. Many fishes have definite spawning grounds far removed from their feeding habitat, and the remarkable migrations of such fishes as the eel and the salmon are directly associated with reproductive instincts (pp. 811 and 861).

The reproductive habits of certain deep-sea fishes are of special interest as an indication of adaptation to the environment. In the lightless,

sparsely populated, abyssal depths it is conceivable that individuals of the sexes may become separated to such a degree that fertilization of eggs at spawning becomes highly fortuitous. In some deep-sea fishes this condition is overcome by the male becoming parasitic on the female, being permanently and securely grown to her body as a mere appendage with a united circulatory system (fig. 231c, p. 831).

The mammals of the sea bring forth living young, which are nursed for a period by the mother. Like those of some fishes, the great migratory movements of whales and seals are associated with wanderings to and from favorable breeding grounds. Growth in whales is extremely rapid; sexual maturity may be reached in two years, and one calf may be produced every other year.

The spawning of many marine animals, especially in boreal waters, is of a spontaneous nature, and vast numbers of individuals spawn within a period of a few days, with the result that in such cases the main spawning period is easily ascertained, since great swarms of eggs or larvae appear suddenly in the plankton and are gradually dispersed by water movements. This feature is especially well illustrated by the oyster, certain sea cucumbers (Cucumaria), nereid worms, and barnacles.

The degree of success of (1) spawning or (2) survival of larvae of successive spawning seasons gives rise to an inequality in numerical strength of year classes of adult or juvenile forms constituting any given population. This inequality is best demonstrated by studies of commercial fishes, investigations of which have been most ardently pursued. However, the same inequality must also occur in the populations of any animals with a normal life span sufficiently long for individuals to live through several reproductive seasons as juveniles and adults.

For purposes of illustration we may consider a species with a life span of several years in which the age of individuals can be accurately determined and in which adequately large and inclusive samples are obtainable for comparison. Now, assume a highly successful spawning and larval survival in a moderate population of this species in the breeding season of 1930, a very poor spawning season in 1931, an average degree of spawning and survival of larvae in 1932, and then another highly successful year in 1933. The 1930 year class will, upon investigation of the whole population in 1931, show up as a disproportionately great number of small, one-year-old individuals in relation to the other age groups in the population. In the next year (1932) the two-year-old individuals of the 1930 spawning are still conspicuous in the population, but the smaller number of one-year-old individuals is evidence of a poor spawning or survival for the 1931 reproductive season. Thus, in 1933 and subsequent years the downward trend of numerical strength of the 1930 and 1931 classes can be traced and compared with other year classes—for example, that of the average year 1932 and of the successful year 1933

(fig. 82). As indicated in fig. 82, the 1930 and 1933 spawning produced “dominant year classes.”

From such comparative studies of year classes and with a knowledge of the spawning habits and age groups, means are provided for analysis of probable environmental factors that determine the degree of success of spawning or survival of larvae, because the relative number of individuals entering into any year class must depend mainly on these critical periods. In subsequent years within the normal life span of the species, the reduction of numbers in year classes is not so likely to be of catastrophic nature. It has been pointed out by Hjort et al (1933) that for a given region the average rate of growth of individuals within the different year classes of Norwegian herring is the same for each year class regardless of the relative numerical strength of the classes. This seems to indicate that, in the sea, nature each year provides a sufficiency of food for the survival and growth of the older stages of this fish as represented in the composite commercial catch. The numbers of individuals belonging to the separate year classes may be widely different (as much as 1 to 30), and this difference must then result from some factor or multiplicity of factors operating to destroy the animals during the very early stages of their existence.


Schematic illustration of changes in year class composition of a population.

Studies of commercially important fish, shellfish, and whales are deeply concerned with analysis of the year classes. For example, in nature there is an equilibrium between the rate at which fish enter the accumulated stock or supply and the rate at which they are withdrawn through natural mortality. Additional removal through fishing exploitation disturbs this equilibrium and may constitute so serious a drain upon the stock that the drain becomes greater than the rate of replenishment and the stock subsequently becomes so depleted that it is no longer

profitably fished. Intensive fishing is evidenced by a decline in the proportion of older and larger specimens entering into the catch and also by a leveling off of abundant year classes. Therefore a record of the trend of the proportion of year classes gives valuable information regarding the toll that fisheries operations are exacting on the accumulated stock in the area being exploited. Such information provides practical aid in formulating conservation policies and in determining the optimum catch practicable to the fishing industry. The theories of fisheries science here involved are of great importance to that branch of marine biology. For a fuller discussion the reader should refer to Hjort et al (1933), Thompson (1937), and other relevant reports.

Year-class analyses of the fish population are also utilized as a basis in arriving at forecasts of the most probable abundance of fish in the next year's catch. Illustrative of this are the investigations of the U. S. Bureau of Fisheries into the fluctuations so characteristic of commercial catches of mackerel. Sette (1931et seq) studied the relative abundance of these fish caught with reference to each year class of the population. The numerical strength of the younger year classes entering into the catch provided a basis for calculating the probable yields that these classes would give in the following year under similar conditions of fishing. Such calculations can be significant only when the downward trend in numbers of the dominant year classes is rather regular. In dealing with migratory fishes, the occurrence of sporadic invasions of populations produced elsewhere or with a different range and whose year-class composition is not known must give rise to unexpected changes in the ratio of the year classes occurring in any one range or locality under investigation. The continued success of such commercial fisheries as the mackerel is determined mainly by the numerical strength of the dominant year classes of the population native to the fishing area. In the above investigation it was found that the big year class of 1923 constituted the main bulk of a declining fisheries yield for a period of years until the industry again experienced a sharp upward incline as the contingents of the 1928 successful spawning entered the catch.


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X. Observations and Collections at Sea



For purposes of oceanographic research, a very sturdy, seaworthy vessel capable of working under practically all weather conditions and of withstanding any storm is required. Vessels engaged in marine investigations can be broadly classified as either oceanic or coastal types, depending largely upon their size and cruising radius, but the two categories are not sharply defined, since large vessels may be used for near-shore investigations and relatively small vessels sometimes extend their operations far out to sea. In the following discussion, vessels and equipment used in coastal surveying and in the study of fisheries problems will not be described, although vessels engaged primarily in such work are sometimes employed in oceanographic investigations. Practically any vessel, small or large, can be used for certain types of investigations, but rarely is any single craft, unless specially designed, suitable for all kinds of oceanographic work. One of the chief requirements of oceanographic vessels operated by private or small organizations is economy of operation. This generally means a relatively small craft with low maintenance cost which can be handled by a small crew. Vessels owned or operated by national agencies, such as the Meteor (Germany), Discovery II (Great Britain), and Willebrord Snellius (Netherlands), are generally fairly large, but in most cases they serve a dual purpose. For example, the Meteor was used as a naval training ship and as a survey vessel, and the Snellius was especially built for surveying work in the Netherlands East Indies.

The following features are desirable in vessels that are to be used in oceanographic research:

  1. Sturdiness and seaworthiness, large cruising radius, and accommodations for laboratory work and the storage of collections.

  2. Low freeboard in order to make possible the handling of instruments near the sea surface and to reduce the wind drift when hove to at stations.

  3. Sails to increase the cruising radius, to provide a safety factor in case of engine breakdown, and to improve the working conditions on board by reducing the roll and vibration when under way. Riding sails to steady the vessel when hove to at stations and to reduce the leeway by keeping the vessel headed into the wind.

  4. Sufficient clear deck space for the installation of winches and for handling bulky equipment such as trawls and dredges.

Name of vessel Nationality Operated by Type of vessel Launched Commissioned for oceanographic work Overall length, feet Tonnage Officers and crew Scientists Reference

aSunk in collision in North Sea, June 23, 1935.

bDestroyed by explosion and fire at Apia, Samoa, Nov. 29, 1929.

cWas previously equipped exclusively for work in terrestrial magnetism and atmospheric electricity.

R.R.S. Discovery II Great Britain Discovery Committee of the Colonial Office Steel steam vessel (trawler) (special) 1929 1930 234 2100 (disp.) 46 6 Ardley and Mackintosh (1936)
F. & V.S. Meteor Germany Hydrographic Department of the Navy Steel steam vessel for survey and training (gunbost) 1915 1924 233 1200 (disp.) 114 10 Spiess (1932a)
H.M.S. Willebrord Snellius Netherlands Hydrographic Section, Department of Defense Steel steam vessel for surveying (gunboat) 1928 1929 204 1055 (disp.) 84 6 Pinke (1938)
R.R.S. Dana (II)[a] Denmark Danis Commission for the Investigation of Sea Steel steam vessel (trawler) 1917 1921 138 360 (gross) 14 8 Schmidt (1929)
Armauer Hansen Norway Geophysical Institute, Bergen Wooden auxiliary ketch (special) 1912 1913 76 57 (gros) 5 6 Helland-Hansen (1914)
Carnegie[b] U.S.A. Department of Terrestrial Magnetism, Carnegie Institution of Washington Wooden auxiliary brigantime (nonmagnetic) 1909 1928[c] 155 568 (disp.) 17 8 Bauer, Peters, Ault, Fleming (1917)
Atlantis U.S.A. Woods Hole Ocean ographic Institution Steel auxiliary ketch (special) 1930 1931 142 460 (disp.) 17 5 Iselin (1933)
E.W. Scripps U.S.A. Scripps Institution of Oceanography, University of California Wooden auxiliary schooner (yacht) 1924 1938 104 140 (disp.) 7 6 Moberg and Lyman (1942)
Catalyst U.S.A. Oceanographic Laboratories, University of Washington Wooden motor vessel (special) 1932 1932 75 94 (gros) 5 9 Thompson (1936)


In table 57 are listed certain representative vessels that have been extensively used in oceanographic investigations. Those owned by national agencies are large, over 200 feet long, and carry large crews, while, on the other hand, vessels owned and operated by institutions are generally between 100 and 150 feet long and carry crews of less than