The Adjacent Seas of the North Atlantic Ocean

The water masses and currents of the North Atlantic Ocean can be dealt with more precisely when the waters of the adjacent seas have been discussed, and we shall therefore turn first to the adjacent seas—the Caribbean Sea and the Gulf of Mexico, the Mediterranean and the Black Seas, the Norwegian Sea, the Baltic and the Polar Seas, and the Labrador Sea and Baffin Bay.

The American Mediterranean Sea. The Caribbean Sea and the Gulf of Mexico can be considered together under the name of the American Mediterranean Sea, a name which has been employed particularly by German oceanographers. The Caribbean Sea has the shape of a wide and somewhat irregular channel, the eastern end of which opens towards the tropical part of the North Atlantic (see fig. 6, p. 35). The opening is wide, but is partly closed by the submarine ridge on which the Lesser Antilles lie. In most localities the ridge rises to less than 1000 m below sea level and communication with the open ocean exists,

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therefore, down to this depth only. The northern boundary of the channel, which is represented by the islands of Cuba, Haiti, and Puerto Rico, is broken by several passages, some of which have greater sill depths than the sill depth of the Lesser Antilles ridge. The important passages are the Anegada and Jungfern Passages to the east of Puerto Rico, and the Windward Passage between Cuba and Hispaniola. The Jamaica Rise between Hispaniola and Honduras divides the Caribbean Sea crosswise in two regions, the eastern and the western Caribbean Seas, of which the latter is also called the Cayman Sea (Parr, 1937). To the northwest the Yucatan Channel opens into the Gulf of Mexico, which has a simpler bottom topography and which, towards the east, is in communication with the Atlantic Ocean through the Straits of Florida. The sill depths in the Yucatan Channel and in the Straits of Florida are approximately 1600 m and 800 m, respectively. The waters of the North Atlantic can therefore pass freely through the American Mediterranean Sea at depths less than about 800 m, but at greater depths communication with the large water bodies of the Atlantic Ocean is more or less restricted.

The water masses of the upper layers enter the Caribbean Sea from the east and show characteristics similar to those of the adjacent North Atlantic waters. Below a thin homogeneous top layer a nearly discontinuous decrease of temperature is found and in most localities a salinity maximum is present within the discontinuity layer. The water masses that enter between the Lesser Antilles show a streaky distribution of salinity, narrow bands of high salinity alternating with bands of low salinity, but these differences are rapidly smoothed out in the Caribbean Sea owing to such intense mixing that the waters flowing through the narrow passages of the Yucatan Channel and the Straits of Florida are of more uniform character.

The upper waters that flow into the Caribbean Sea are mainly of North Atlantic origin but contain a considerable admixture of South Atlantic water. The relative amounts of the two types of water can be approximately determined by a comparison between the temperatures and salinities of the South Atlantic water that crosses the Equator, the waters passing through the Yucatan Channel, and the water from the western Sargasso Sea, the T-S relations of which have been examined by Iselin (1936). Table 77 contains corresponding temperatures and salinities of these three water masses at different σ_t values. The last column in the table gives the ratio between the amounts of South Atlantic and North Atlantic water that pass through the Yucatan Channel. It is seen that at higher temperatures the current through the Channel carries about one part of South Atlantic and three and one half parts of North Atlantic water. The total transport is about 26 million m³/sec,

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of which, accordingly, about 6 million m³/sec represent South Atlantic water.

Below the upper water a considerable amount of Antarctic Intermediate Water enters the Caribbean Sea, such that the water at the intermediate salinity minimum is mainly of South Atlantic origin, as pointed out by Nielsen (1925) and Parr (1937). The intermediate salinity minimum in the Caribbean Sea decreases in intensity in the direction of flow. In the eastern part of the Caribbean Sea, in about longitude 68°W, the average minimum salinity is 34.73 ‰, whereas in the Yucatan Channel the average minimum salinity is 34.88 ‰. This increase in salinity is, according to Seiwell (1938), mainly due to vertical mixing within the moving water mass.

The deeper portion of the American Mediterranean Sea is divided into a series of basins, the most important being the Venezuela and Colombia basins in the eastern Caribbean Sea, the Cayman Trough and the Yucatan Basin in the western Caribbean Sea, and the Mexico Basin in the Gulf of Mexico. The character of the deep water in these basins depends upon their communications with the adjacent parts of the North Atlantic Ocean and upon the intercommunications between the basins. The greatest depth of the sills separating the basins of the eastern Caribbean from the North Atlantic are found in the Anegada and Jungfern Passages, between which lies the small St. Croix Basin. Figure 173 shows the distribution of potential temperature in a section passing from the Atlantic through the Anegada Passage, the St. Croix Basin, and the Jungfern Passage into the Venezuela Basin (Dietrich, 1939). It appears from this distribution that the sill depth of the Anegada Passage is somewhat greater than that of the Jungfern Passage,

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and that the waters in the St. Croix Basin therefore show a lower potential temperature, about 3.5°, as compared to 3.85° in the Venezuela Basin. The deep water in the Venezuela Basin is of lower density than that of the adjacent Atlantic Ocean, for which reason the deep water of the basin is renewed by inflow across the sill. The basin therefore belongs to the second type (see p. 148). The renewal of the deep water must be a relatively rapid process, because the oxygen content is high and has, according to Seiwell (1938), an average value of approximately 5.1 ml/L at a depth of 2500 m.

The sill depth of the Jamaica Rise, which separates the western Caribbean Sea from the eastern, is less than the sill depth of the Windward Passage. The deep water of Cayman Trough and the Yucatan Basin is therefore renewed through the Windward Passage which, according to the distribution of potential temperature (fig. 173), must have a sill depth of approximately 1600 m. The renewal of the deep water in these basins appears to be even more rapid because the average oxygen content in the Yucatan Basin and the Cayman Trough is, at 2500 m, between 5.5 and 6.0 ml/L.

The deepest connection between the Mexico Basin and the adjacent seas is found in the Yucatan Channel, and the deep water of the Mexico Basin is therefore renewed from the Yucatan Basin by flow across the sill in the Yucatan Channel. The sill depth is not exactly known but is probably somewhat less than the sill depth of the Windward Passage, because the potential temperature of the water in the Mexico Basin is about 3.95° as compared to 3.85° in the Yucatan Basin. The passages to the north of Cuba are much more shallow and no renewal of deep water can take place through them. Even in the case of the Mexico

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Basin, renewal must be fairly rapid because the oxygen content at 2500 m averages about 5.0 ml/L.

The different sill depths have not yet been established by means of soundings, and the depths given above are all based on studies of the hydrographic conditions. These can be interpreted in several manners and different values of the sill depths have therefore been arrived at by different authors, but in the above we have followed Dietrich's presentation (1939), which appears to combine the known facts in the most satisfactory manner. Dietrich has also been able to plot the potential temperatures at depths greater than 2500 m, and these show a remarkable regularity in spite of the fact that the greatest differences observed in the single basins do not exceed 0.05°. The arrangement of the potential temperatures further supports the conclusions as to the localities at which inflow of deep water into the basins takes place.

The salinity of the deep water is very uniform. According to Parr (1937) the average salinity is 34.98 ‰ and the deviations from this value are within the probable errors of the salinity determinations.

It is evident from this summary that the American Mediterranean Sea does not exercise any influence upon the deep-water circulation of the Atlantic Ocean, but it may exercise a considerable influence on the circulation in the upper layers.

The surface currents in spring are shown in fig. 174. A strong current passes through the Caribbean Sea, continues with increased speed through the Yucatan Channel, bends sharply to the right, and flows with great velocity out through the Straits of Florida. On the flanks of the main current numerous eddies are present, of which the one in the wide bay between Nicaragua and Colombia and the one between Cuba and Jamaica are particularly conspicuous. In the Gulf of Mexico several large eddies exist, and all of these appear to be semipermanent features, the locations of which are determined by the contours of the coast and the configuration of the bottom.

The presentation of the currents in fig. 174 is based on ships’ observations. When attempting a calculation of currents by means of numerous Atlantis data from the Caribbean Sea, Parr (1937) found that the flow is not directed parallel to the contours of the isobaric surfaces, but between the Lesser Antilles and the Yucatan Channel the current flows uphill. Parr suggested that this feature may be due to piling up of water in front of the narrow Yucatan Channel, caused by the stress exerted on the surface by the prevailing easterly winds. This idea was further examined by Sverdrup (1939b), who concluded that the piling up of the surface water can be fully explained as the effect of winds blowing with an average velocity of about 10 m/sec, which agrees well with the observed values in spring. A further consequence of this piling up is that in the Gulf of Mexico a higher sea level is maintained than along the adjacent coast of

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the United States facing the Atlantic Ocean. At Cedar Keys, off the southwest tip of Florida, the average sea level is 19 cm higher than the average sea level at St. Augustine, Florida, on the east coast. This indicates that the prevailing winds over the Caribbean Sea produce a hydrostatic head that may, according to Montgomery (1938), account for the major part of the energy of the Florida Current (p. 673).

Regardless of the effect of the stress of the wind or the existence of a hydrostatic head, the mass distribution must adjust itself to the major currents that are present. In general, the lighter water is therefore found on the right-hand side of the current and the sea surface rises to the right when looking in the direction of flow. This rise to the right is particularly conspicuous within a swift and narrow current and, in the case of the Florida Current, it has been computed that at the north coast of Cuba sea level must be about 45 cm higher than at Key West, Florida (Dietrich, 1936, Montgomery, 1938). The Florida Current will be further dealt with when discussing the Gulf Stream System of the North Atlantic, but in this place it should be pointed out that, according to Dietrich (1939), the Florida Current is essentially a direct continuation of the current through the Yucatan Channel and that the waters of the Gulf of Mexico mainly form independent eddies and are only to a small extent drawn into the Straits of Florida.

The European Mediterranean Sea. From the oceanographic point of view the waters of the Mediterranean Sea are basin water masses which are in communication with the Atlantic Ocean only through the

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narrow and shallow Strait of Gibraltar, the exchange of water through the Suez Canal being negligible (p. 688). The Mediterranean body of water can again be divided into several smaller ones, each of which has its specific characteristics, the most outstanding being the waters of the Black Sea, which are in restricted communication with the Mediterranean Sea proper through straits at the Dardanelles and the Bosporus. Geographically the Black Sea is not considered a part of the Mediterranean Sea, but oceanographically it should be regarded as such.

The Mediterranean proper is divided into a series of deep basins more or less isolated from each other. The most outstanding of these are, in the western Mediterranean, the Algiers-Provençal Basin to the west of Sardinia and Corsica and the Tyrrhenian Basin on the west side of Italy, and, in the eastern Mediterranean, the Ionian Basin to the south of Italy and Greece and the Levantine Basin to the south of Asia Minor (fig. 5, p. 34). The two former basins are separated from the two latter by shallow ridges between Tunisia and Sicily, whereas the two basins in the western Mediterranean Sea and the two basins in the eastern Mediterranean Sea are separated by deep sills.

In the Mediterranean Sea proper (not including the Black Sea), evaporation greatly exceeds precipitation and runoff, and in winter deep water of very high salinity is formed in different places by vertical convection currents. The Mediterranean proper belongs, therefore, to the first type of basin (p. 147), in which deep water of great density is formed with outflow over the sill and inflow at the surface.

The water masses of the Mediterranean are characteristically different from those of the adjacent North Atlantic, because, owing to the isolation, the deep water has a higher temperature. The Atlantic water which flows into the Mediterranean Sea is so rapidly mixed with Mediterranean surface water that it soon loses its Atlantic character. The surface salinity is higher than 37.00 ‰, except where the surface current flows to the east along the north coast of Africa as far east as Tunisia, and in the inner portion of the Adriatic Sea and the Aegean Sea, where there is a considerable addition of river water or surface water from the Black Sea. It is above 39 ‰ to the south of Asia Minor. The surface salinity is subjected to seasonal variations, but these are not known in detail. The surface temperature generally increases from the Strait of Gibraltar to the inner portions, except in winter, when the lowest surface temperatures are found in the most northerly portions, namely off the French and Italian Rivieras and in the inner portions of the Adriatic and Aegean Seas. The seasonal variation in temperature is great, the annual range everywhere exceeding 9° and reaching 13° to 14° off the Riviera and in the northern part of the Adriatic Sea.

Below the surface four different water masses are encountered, including a thick transition layer which separates the intermediate water from

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the deep water. The character of these water masses is illustrated in fig. 175, showing the vertical distribution of temperature and salinity at two Dana stations, station 4119 in the Tyrrhenian Sea and station 4070 in the Ionian Sea.

The surface water generally extends to a depth of 100 to 200 m. In the western Mediterranean the lower limit of the surface water is indicated by a temperature minimum, as seen at station 4119, whereas in the eastern Mediterranean the temperature minimum is generally absent and a layer of slow temperature decrease is found instead, as seen at station 4070. The depth of the temperature minimum or the depth of the layer of slow decrease represents the depth to which vertical convection currents ordinarily reach in winter, and the temperature and salinity at the depth of the minimum or in the layer of small decrease correspond to the surface values in winter.

Below the surface layer one finds, in the greater part of the Mediterranean, an intermediate water which is characterized by a salinity maximum at a depth of 300 to 400 m. In the western sea a temperature maximum is present at nearly the same depth, but in the eastern sea only a second layer of small temperature decrease is found. This intermediate water, according to Nielsen (1912), is formed in the inner part of the Mediterranean, from where it spreads towards the west and finally flows out through the Strait of Gibraltar. At the Thor stations 178 and 160, in about long. 30°E, no subsurface salinity maximum was present but the salinity decreased regularly from the surface. Between 175 and 200 m the temperature remained constant, and this layer of constant temperature represents both the layer to which convection currents reach in winter and the layer of intermediate water which originates in that part of the sea. As this water spreads to the west, the salinity and the temperature at the core decrease owing to mixing processes.

The lower boundary of the intermediate water can be placed, somewhat arbitrarily, at about 600 m, where both temperature and salinity

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decrease rapidly. This decrease continues within the transition layer between the intermediate water and the deep and bottom water which, in the Mediterranean basins, is found below a depth of 1500 to 2000 m. At this depth a temperature minimum is encountered, and at greater depth the temperature increases adiabatically towards the bottom, as first shown by Nielsen. The potential temperatures at stations 4119 and 4070, at depths greater than 1500 m, are also plotted in fig. 175, from which it is seen that at station 4119 the potential temperature between 2000 and 3400 m was constant at 12.74° and at station 4070 the potential temperature between 2500 and 4200 m was constant at 13.19°. The salinity appeared to decrease slightly towards the bottom, but the decrease was nearly within the limits of accuracy of the observations. At station 4119 the average salinity below 2000 m was 38.42 ‰ and at station 4070 the average salinity below 2500 m was 38.63 ‰, the greatest deviation from the averages being 0.02 ‰. Therefore, the Mediterranean basins are filled by water of uniform character, although the type varies slightly from one basin to another. According to the Thor and Dana observations the average values of temperature and salinity at 2000 m are:

These data show that there is an appreciable difference between the western and eastern basins, which are separated by the rise between Tunisia and Sicily.

The deep and bottom water in the different basins is formed in localities in which the intermediate water is lacking and where convection currents in winter can reach from the surface to the bottom. Nielsen shows that the bottom water of the western basins originates mainly from the northern parts of the Balearic and Ligurian Seas, off the French and Italian Rivieras. At the Thor station 37, which was occupied on January 30, 1909, in lat. 41°56′N, long. 6°18′E, nearly uniform density was found between the surface and 2000 m. The average value of σ_t was 29.04, and the greatest deviations from the average were ±0.04. The temperature at the surface was 12.40°, and the potential temperature at 2000 m was 12.70°, the salinities were 38.24 ‰ and 38.35 ‰, respectively. Even more extreme conditions were observed by Nathansohn

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(Krümmel, 1911) off Monaco on April 7, 1909, when the temperature at the surface was 12.67°, the potential temperature at 2213 was 12.62°, and the salinities were 38.51 ‰ and 38.49 ‰, respectively. Probably no bottom water is formed in the Tyrrhenian Sea, and the slightly higher temperatures and salinities of the deep water are due to a small admixture of intermediate water. The deep and bottom water of the eastern basins is formed, according to Nielsen, in the Aegean Sea and in the southern part of the Adriatic Sea, from where it spreads to the south and west.

Nielsen believes that the formation of bottom water takes place intermittently and this concept is substantiated by the potential temperatures at station 4070, which were constant at 13.38° between 1500 and 2000 m and at 13.19° below 2500 m, indicating that slightly different water is formed under different conditions (fig. 175). Nielsen furthermore emphasizes that the uniform character of the deep and bottom water in a horizontal direction demonstrates active horizontal movement at great depth.

The water that sinks from the surface in order to renew the deep water must be replaced, and such replacement is accounted for by an inflow of Atlantic water. The exchange of water between the North Atlantic Ocean and the Mediterranean takes place through the Strait of Gibraltar, which is noted for its strong currents. The bottom topography of the Strait is complicated; in its shallow portion two or possibly three ridges are present with sill depths of about 320 m. The character of the water masses passing in and out through the Strait of Gibraltar is illustrated in fig. 176, showing isohalines and isotherms in a longitudinal section through the Strait. The observations used for perparing this section were made in late spring and early summer, in the months of May, June, and July of different years (Schott, 1928, Ramalho and Dentinho, 1931).

In the upper layers Atlantic water of a salinity somewhat higher than 36.00 ‰ and a temperature

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higher than 13° flows in, whereas water of a salinity higher than 37.00 ‰ and a temperature about 13° flows out along the bottom. The deeper water in the Mediterranean Sea inside the Strait of Gibraltar has a salinity of about 38.40 ‰ and a temperature of 13°, but in the Strait intensive mixing takes place, whereby the salinity of the outflowing water is greatly reduced and that of the inflowing is increased. In the Strait the inflowing water has a thickness of approximately 125 m, but the boundary surface separating the in- and outflowing layers lies deeper on the African side of the Strait and the greater inflow therefore takes place on the southern side. This inclination of the boundary surface is due to the effect of the earth's rotation. The average velocity at the surface in some localities is in excess of 200 cm/sec (4 knots), and close to the African coast a narrow counter-current with velocities up to 100 cm/sec (2 knots) is often encountered. Superimposed on the currents carrying water in and out of the Mediterranean Sea are strong tidal currents which greatly reduce the inflow when they are directed from the Mediterranean Sea to the Atlantic Ocean, and greatly increase the inflow when they are directed from the Atlantic to the Mediterranean. The inclination of the boundary surface probably varies with the speed of the inflowing current, and great vertical oscillations of tidal period therefore take place.

The average velocity of the total inflow, according to measurements on Danish expeditions, is approximately 100 cm/sec (2 knots). By means of this value Schott (1915) has computed the inflow to be approximately 1.75 million m³/sec. The outflow can be computed by means of the relations on page 148. Such computations have been made by Nielsen (1912) and Schott (1915), but subsequent data indicate that they have assumed somewhat too high values for the salinity of the outflowing water. From the sections published by Schott in 1928 it appears that the average salinity of the inflowing water should be, in the Strait of Gibraltar, about 36.25 ‰ and that of the outflowing water should be not more than 37.75 ‰. With these values one obtains an outflow of 1.68 million m³/sec and the difference between inflow and outflow, 70,000 m³/sec, represents the excess of evaporation over precipitation and runoff. A fraction of this excess is made up, however, by a net inflow from the Black Sea, amounting to 6500 m³/sec (p. 650).

The exchange of water through the Strait of Gibraltar presents a good example of how water from one region can be transformed by external influences and return as a different type of water, in this case as water of high salinity. The rapidity of the exchange is illustrated by stating that the in- and outflow is sufficient to provide for a complete renewal of the Mediterranean water in about seventy-five years.

Using the figures which Schott gives for precipitation and runoff, one arrives at the water budget of the Mediterranean Sea proper, which is summarized in table 78. From this it appears that the total evaporation amounts to 115,400 m³/sec, corresponding to an annual evaporation of 145 cm, which is in fair agreement with the annual evaporation in these latitudes according to observations and computations (fig. 27, p.121).

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As one might expect, the evaporation from the Mediterranean Sea is somewhat greater than that from the open ocean in the same latitude, which is 110 cm a year.

This discussion brings out a very characteristic difference between the Mediterranean Sea and the American Mediterranean Sea. The Mediterranean Sea exercises no great influence on the surface currents of the Atlantic Ocean because the amount of water that flows in through the Strait of Gibraltar represents a small fraction of the water masses that are transported by currents of the upper water layers; but the Mediterranean exercises a widespread influence on the deep water of the North Atlantic by adding appreciable quantities of water of high salinity. The American Mediterranean Sea, on the other hand, exercises a great influence upon the currents of the upper layers because large water masses flow into the Caribbean Sea and out of the Gulf of Mexico, and the types of currents are greatly modified by the character of the passages; but the importance of the American Mediterranean Sea to the deep-water circulation is negligible.

Within the Mediterranean an exchange of water takes place across the submarine ridge between Tunisia and Sicily that is similar to the exchange through the Strait of Gibraltar. Surface water of Atlantic origin and of salinity about 37.20 ‰ flows towards the west through the channel which cuts the ridge between Tunis and Sicily at a depth of about 400 m. According to Nielsen about 4 per cent of the inflowing water is lost by excess evaporation in the eastern Mediterranean, whereas 96 per cent is carried out again by the intermediate current after the salinity has been increased by about 1.5 parts per mille.

In fig. 177 a schematic picture is given of the surface currents and the flow of the intermediate water (according to Nielsen, 1912). It is seen that both surface currents and intermediate currents have a tendency to circle the different areas in a counterclockwise direction.

The distribution of oxygen confirms the conclusions as to the origin of the different water masses of the Mediterranean and the flow of water.

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The surface water, which extends to a depth of 100 to 200 m, shows a high oxygen content. The oxygen content of the intermediate water is relatively high in the eastern Mediterranean where this water is formed, but decreases toward the west, the lowest values being found in the western part of the Mediterranean area, the Balearic Sea. The transition layer between the intermediate and the deep water is characterized by an oxygen minimum that is more conspicuous in the eastern than in the western regions. The deep water has a somewhat higher oxygen content, the highest value being found in the Balearic Sea, where the formation of deep water is probably most rapid. These general features are shown in table 79, which is based upon the Thor and the Dana observations in the summers of 1910 and 1930. For three different areas, the Ionian, Tyrrhenian, and the Balearic Seas, the table contains mean values of temperature, salinity, and oxygen at the layer of salinity maximum, at the layer of oxygen minimum, and below 2000 m. It also gives average depths of the salinity maximum and the oxygen minimum within the different areas. The observations from 1910 and 1930 agree remarkably well in the western Mediterranean, but in the Ionian Sea the oxygen values and the salinity values were on the whole higher in 1930 than in 1910, indicating that in different years the formation of intermediate water may be more or less intense.

The water masses of the Black Sea are entirely different from those of the Mediterranean proper. In the Black Sea precipitation and runoff exceed evaporation, for which reason a surface layer is present of relatively low salinity and correspondingly low density. Vertical mixing tends to reduce the density of the deep water, which therefore has a lower density than that of the water in the Mediterranean Sea proper at the same depth. The character of the waters of the Black Sea is seen from table 80, containing some of the observations at Thor Station 172, about 100 m inside the Bosporus.

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Owing to the distribution of density, the surface waters of the Black Sea flow out through the Bosporus and through the Dardanelles, and Mediterranean water flows in along the bottom. Intensive mixing takes place in these narrow straits, and the salinity of the inflowing water is therefore reduced from more than 38.50 ‰ at the entrance of the Dardanelles to between 35.00 and 30.00 ‰ where the bottom current enters the Black Sea at the northern end of the Bosporus. Similarly, the salinity of the outflowing water is increased from about 16.00 ‰ where it enters the Bosporus to nearly 30.00 ‰ where it leaves the Dardanelles. The out- and inflowing water masses are separated by a well-defined layer of transition which oscillates up and down according to the contours of the bottom. The currents through the Bosporus can well be compared to two rivers flowing one above the other in a river bed which has a width of about 4 km (2 miles) and a depth of 40 to 90 m.

Appreciable water masses pass through the Bosporus. According to current measurements and other observations by A. Merz (L. Möller, 1928), the most probable values of the out- and inflow through it are 12,600 m³/sec and 6,100 m³/sec, respectively. (The Mississippi River carries, on an average, 120,000 m³/sec.) The difference between inflow and outflow, 6,500 m³/sec, represents the excess of precipitation and runoff over the evaporation. Precipitation and runoff have been estimated at 7,600 and 10,400 m³/sec, respectively, and the evaporation should therefore amount to 11,500 m³/sec or 354 km³/year. The area of the Black Sea is 420,000 km² and the above value therefore corresponds to an evaporation of 84 cm per year, in good agreement with observed and computed values for this latitude (p. 121).

For the Black Sea the out- and inflowing water masses are of the same order of magnitude as the difference between precipitation plus

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run-off and evaporation, whereas for the Mediterranean proper the in- and outflowing water masses are many times greater than the excess of evaporation. This is in agreement with the general considerations which were set forth on page 147. It follows that the renewal of the deep water of the Black Sea is a very slow process because the entire renewal has to be taken care of by the water which flows in along the bottom of the Bosporus. This inflow is so small in proportion to the total volume of water in the Black Sea that complete renewal of the water below a depth of 30 meters would take about 2500 years. For all practical purposes the deep water of the Black Sea is stagnant, as is evident from the fact that below a depth of 200 m the water contains no oxygen, but large quantities of hydrogen sulphide (table 80).

The Arctic Mediterranean Sea. Under the name “Arctic Mediterranean Sea” are understood the areas to the north of the Wyville Thomson Ridge between the Orkneys and Faeroe Islands, the Faeroe Islands-Iceland Ridge, and the Iceland-Greenland Ridge. Exchange of water between the Atlantic Ocean and the Arctic Mediterranean Sea takes place across these ridges. In addition, the water masses of the Arctic Mediterranean are in restricted communication with those of the Atlantic through the English Channel and through the narrow sounds to the west of Greenland, and with the waters of the Pacific Ocean through Bering Strait. The Arctic Mediterranean is the largest of all adjacent seas and is subdivided into a number of distinct areas, of which the more important are the Norwegian Sea, the North Sea, the Baltic

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with its gulfs, the Barents Sea, and the North Polar Sea. The areas of some of these parts are shown in table 4 (p. 15).

Much the major exchange of water between the Arctic Mediterranean and the adjacent oceanic regions takes place through the straits between Scotland and Greenland. The character of the exchange is radically different from that between the Atlantic Ocean and the Mediterranean Sea, where inflow of Atlantic water takes place in the upper layers of the Strait of Gibraltar and outflow near the bottom. The opening between the Atlantic Ocean and the Arctic Mediterranean is so wide that the water flows in through the southeastern part of the opening and out through the northwestern. The character of the water masses is illustrated

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in fig. 178. The top part of the figure contains a chart of the area between northern Scotland and eastern Greenland on which the 1000-m bottom contour is shown. On the chart are indicated the location of two sections, one through the Faeroe-Shetland Channel to the north of the Wyville Thomson Ridge, and one through the Denmark Strait between Iceland and Greenland to the north of the Iceland-Greenland Ridge (fig. 3, p. 30). In the lower part of the figure are shown the distributions of temperature and salinity in these sections according to observations of the Michael Sars in 1910 (Helland-Hansen, 1930) and observations on board the Heimdal in 1933 (Helland-Hansen, 1936). The principal inflow takes place across the Wyville Thomson Ridge. The inflowing water masses are characterized by a salinity higher than 35.00 ‰ and a temperature higher than 4°, the larger part of the water masses having a salinity above 35.25 ‰ and a temperature above 8°. The section through Denmark Strait shows that a small amount of Atlantic water of salinity above 35.00 ‰ and temperature above 4° flows north along the west coast of Iceland, but by far the greater amount of the water masses shown in the left-hand sections of fig. 178 flow out, having temperatures between 2° and -1.5°, and salinities ranging from 34.90 ‰ to 31.00 ‰.

The great difference between the inflowing Atlantic water masses and the outflowing Arctic water is illustrated in fig. 179. In the main part of the figure, T-S curves are plotted showing the character of the waters at Michael Sars station 106 (A), which was occupied in the Faeroe-Shetland Channel on August 10 and 11, 1910, and at the Heimdal station 47 (B), which was occupied in the channel to the north of the Iceland-Greenland Ridge on July 22, 1933. The inset diagram shows T-S curves which are drawn on a much wider salinity scale, demonstrating the character of the waters at Michael Sars station 106 (A); at Michael Sars station 100 (C), which was occupied to the southwest of the Wyville Thomson Ridge on August 6, 1910; and at the Heimdal station 33 (D), which was occupied to the south of the Iceland-Greenland Ridge on July 17, 1933. The last two stations have been added in order to demonstrate the similarity between the waters in the Faeroe-Shetland Channel above 600 m and the water masses to the south of the Ridge on the entire distance from the Orkneys to Greenland. The inset diagram demonstrates that one has to deal here with exactly the same water mass, but water which to the north of the Wyville Thomson Ridge was found at a depth of 600 m, south of the ridge was present at 1200 m, and to the west of Iceland was present at 800 m.

The similarity of the water masses at Michael Sars station 100 and Heimdal station 33 shows that only part of the Atlantic water flows across the Wyville Thomson Ridge and that part bends towards the west and flows towards Greenland. This current will be dealt with later

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(p. 682), but here it will be emphasized that the Orkney-Greenland Ridge forms a most effective barrier between the deeper waters of the Atlantic and those of the Arctic Mediterranean.

On the southern side of the barrier, Atlantic water is present to the greatest depths except along the continental shelf of Greenland, where the Polar Current flows south, but on the northern side of the barrier and below the sill depths the water masses have a uniform salinity of about 34.92 ‰ and a temperature below 0°. This deep water, formed in the Norwegian Sea, has a greater density than the Atlantic water at corresponding depths, but there is no evidence of an outflow of this water across the sills. At the Wyville Thomson Ridge the deep water of the Norwegian Sea does not reach up to the sill depth, whereas at the Iceland-Greenland Ridge it does, as is evident from the Heimdal observations, but no trace of that deep water was present at Heimdal station 33 on the southern side of the ridge (see curve D, fig. 179) nor at any of the Meteor stations in the same region (Defant et al, 1936).

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Helland-Hansen (1934) has computed the volumes of Atlantic water that in May, 1927 and May, 1929, flowed north through a section from lat. 60°40′ on the coast of Norway to lat. 63°20′, long. 4°W, northwest of the Faeroe Islands. He found 2.78 and 3.31 million m³/sec respectively, and the average inflow of Atlantic water may therefore be estimated at about 3 million m³/sec.

In addition to the inflow of Atlantic water, an inflow of Bering Sea water takes place through Bering Strait. On August 1, 1934, the United States Coast Guard Cutter Chelan anchored in Bering Strait and measured currents between the surface and the bottom during 21 hours. From these measurements it was concluded (U. S. Coast Guard, 1936) that the transport inward through Bering Strait was 0.88 million m³/sec, but the average annual transport is probably much smaller because the inflow is less regular and weaker in winter. It does not seem probable that the average inflow during the year exceeds 0.3 million m³/sec.

The Arctic Mediterranean also receives a considerable amount of fresh water in the form of runoff from the great Siberian and Canadian rivers and from excess of precipitation over evaporation. The runoff from the Siberian rivers averages, according to Zubov (1940), about 0.16 million m³/sec. The excess precipitation over the North Sea and the Norwegian Sea can be estimated at about 0.5 m per year, and over the Arctic Sea, where the precipitation is small, at about 0.12 m per year. On this basis one finds a total excess of precipitation of about 0.09 million m³/sec.

The outflow from the Arctic Mediterranean Sea takes place principally through the Denmark Strait, for which reason the water budget of the region can be presented as follows:

On the basis of these figures one finds that a complete renewal of the waters of the Arctic Mediterranean Sea would take about 165 years.

A rough check on the relative correctness of the above values can be obtained by considering that the net salt transport into the Arctic Mediterranean must be zero. The average salinity of the inflowing Atlantic water is about 35.30 ‰ and that of the Bering Sea water about 32.00 ‰. Therefore 3.0 × 35.3 + 0.3 × 32.0 = 3.55 S_A where S_A is the salinity of the outflowing Arctic water. This equation gives S_A = 32.5 ‰, in fair agreement with the salinity shown in the section in fig. 178. If the addition of fresh water were 0.1 × 10⁶ m³/sec or 0.4 × 10⁶ m/sec, the salinity of the outflowing water would be 34.0 ‰

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or 31.2 ‰, respectively, but of these values the first appears to be too high and the second appears to be too low, whereas the assumed values of 0.25 × 10⁶ m³/sec leads to a reasonable result.

The amount of heat given off by the water can also be estimated. The average temperatures of the in- and outflowing waters between Scotland and Greenland can be taken as 8°C and -1°C, respectively, and the average temperature of the water flowing through Bering Strait, the runoff, and the excess precipitation can be taken as 0°C. The total amount of heat given off by the waters in the Arctic Mediterranean Sea is then about 24 × 10¹² g cal/sec. It is probable that at least half of this amount is given off where the Atlantic water flows north along the west coast of Norway, or over an area not greater than 2 × 10¹² m². The average amount of heat given off in this area would then be 12 g cal/m²/sec = 0.072 g cal/cm²/min = 103 g cal/c m²/day. This value, although uncertain, serves to demonstrate the important bearing of the Atlantic Current on the climate of the extreme northwestern part of Europe.

From the above discussion of the water, salt, and heat budgets of the Arctic Mediterranean Sea it is evident that the Atlantic water which contributes mostly to the inflow is diluted and cooled off to such an extent that the outflowing water is of entirely different character. We have here another striking example of local factors operating towards changing the inflowing water and producing a new water mass.

The influence of the local factors is more or less conspicuous in the different parts of the Arctic Mediterranean Sea, giving rise to other water masses than those mentioned, and to water masses produced by mixing. Every subdivision of the Arctic Mediterranean has its own characteristic water masses, which will be briefly described.

In the Norwegian Sea, Atlantic water is found off the west coast of Norway, where it flows to the north, losing some of its heat content to the atmosphere and being somewhat diluted by excess precipitation. On the right-hand side of the Atlantic water is the Norwegian coastal water, which has a lower salinity, owing to runoff, and a considerable annual range in surface salinity and temperature. On the left-hand side of the Atlantic water are found water masses which have been formed by mixing between the Atlantic water and the Arctic water which flows south along eastern Greenland. The latter is characterized by low salinity and temperatures below 0°C, as illustrated in the left-hand section in fig. 178. The mixed water in the central and western parts of the Norwegian Sea has a salinity around 34.90 ‰ and at the surface a temperature which varies considerably during the year. In winter the surface layers are cooled, but before reaching freezing point the waters attain a higher density than that of the deeper waters and therefore sink to the bottom. By this process, first described by Nansen (1906), the bottom

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water of the Norwegian Sea is renewed. As further evidence for the correctness of the explanation, Helland-Hansen and Nansen (1909) point out that surface samples, taken by sealing vessels to the northeast of Jan Mayen in March to May, show temperatures between -1.2° and -1.9°, and salinities between 34.70 ‰ and 34.94 ‰. The bottom water has a salinity between 34.92 ‰ and 34.93 ‰ and in the northern part a temperature of about -1.1° to -1.2°C, while in the southern part the temperature is about -1.0°C. This uniform bottom water fills all the basins of the Norwegian Sea at depths below 600 m, but above 1500 m the temperatures are somewhat higher than those mentioned.

The North Sea waters have, in general, salinities between 34.00 ‰ and 35.00 ‰, but Atlantic water of salinity above 35.00 ‰ is found in a tonguelike area to the south of a line from Scotland to the west coast of Norway and in another tonguelike area extending northwest from the English Channel. Norwegian coastal water of salinity as low as 30.00 ‰ is encountered in the northeastern part of the North Sea (Deutsche Seewarte, 1927).

The Baltic Sea, together with the Gulf of Finland and Gulf of Bothnia, represents a region which, oceanographically, has some features in common with the Black Sea. In the Baltic, as in the Black Sea, precipitation and runoff greatly exceed evaporation, for which reason a layer of brackish surface water is formed that flows out through the sounds between Sweden, the Danish Islands, and Jutland. Along the bottoms of these sounds, North Sea water that has been considerably diluted with Baltic water flows in. The sounds are narrow and shallow, for which reason intensive mixing takes place in and directly outside the sounds, and the influence of the Baltic water does not reach to any great distance. In this respect conditions are similar to those in the Black Sea and the Mediterranean, where the influence of the low-salinity water flowing out from the Black Sea is present only in the northern part of the Aegean Sea. There exists, however, one striking difference. The Black Sea has a deep basin filled by stagnant water of relatively high salinity, but in the Baltic only a few small basins exist, the deepest (with a maximum depth of about 210 m) being found off the island of Gotland. In these basins water of higher salinity is found, but the basins are so limited in extent and the difference in salinity between the bottom water and the surface water is so small that renewal of the deep water in the basins can take place, and stagnant water corresponding to that of the Black Sea is not found. Over large portions of the Baltic the salinity of the surface layer is about 7.00 ‰ but it drops, in the inner parts of the Gulfs of Finland and Bothnia, to less than 3.00 ‰, in spring even below 1.00 ‰ (Deutsche Seewarte, 1927). The salinity of the water in the different basins decreases from about 16.00 ‰ in the southern basins to about 12.00 ‰ in the basins further north. The surface temperatures show a large annual variation.

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In the Gulfs of Finland and Bothnia ice forms during winter, and these gulfs remain ice-covered for periods between three and five months, depending upon the severity of the winter. In very severe winters the sounds between Sweden, the Danish Islands, and the Danish mainland also freeze over.

In the North Polar Sea three main water masses are encountered, the Arctic Surface Water, the Atlantic water, and the Arctic Deep Water. The Arctic Surface Water has a low salinity ranging from a few per mille at the mouths of the Siberian rivers to 32 ‰ or 33 ‰ to the north of Spitsbergen. Owing to the low salinity of the surface waters, no deep water is formed in the Arctic Sea itself, but in winter a top layer of homogeneous water is developed. On the North Siberian Shelf at a distance of about 400 km from the coast, the salinity of this homogeneous layer reached, in 1923, its maximum value, 29.67 ‰, in May (Sverdrup, 1929), but at the end of the summer a nearly fresh top layer was formed by the melting of ice, and as a result of admixture of this fresh water the salinity was somewhat reduced down to a depth of about 30 m. The temperature of this upper layer remained at freezing point during the winter and in summer was raised slightly above freezing point. It is probable that an upper layer of similar characteristics is found all over the Polar Sea, as indicated by Nansen's observations in 1893–1896, but data from later Russian expeditions are not yet available.

Below this upper layer and a transition layer is found the Atlantic water, which enters the North Polar Sea north of Spitsbergen as a subsurface flow. In 1931 (Mosby, 1938) water of a salinity of 35.10 ‰ and a temperature between 3° and 4° was found to the north of Spitsbergen in lat. 80°38′ and longitude 13°41′E, between depths of 75 and 400 m. Nansen's observations (1902) showed that this Atlantic water can be traced across the Polar Sea toward the region of the New Siberian Islands, and his observations have recently been confirmed by those which were conducted during the drift of the Sedov in 1937–1940. Zubov (1940) states that both the temperature and the salinity of the Atlantic water were higher in 1937–1940 than they were in 1893–1896. In both instances admixture with Polar water led to a decrease of temperature and salinity in the direction in which the Atlantic water spreads out.

The deep water of the Polar Sea has a uniform salinity of about 34.93 ‰ and uniform temperature of about -0.85°. Nansen observed an increase of the temperature toward the bottom which can be explained as a result of adiabatic heating. This deep water cannot be formed anywhere in the Polar Sea, but it shows great similarity to the deep water in the northern part of the Norwegian Sea. Nansen therefore concluded that the deep water of the Polar Sea was formed in the Norwegian Sea and flowed in across the submarine ridge between Spitsbergen and Greenland, which supposedly has a sill depth of 1200 to 1500 m. These conclusions

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were confirmed by the observations made on board the Nautilus (Sverdrup, 1933). It is not yet known whether the observations on recent Russian expeditions will modify the earlier results.

The marginal areas of the North Polar Sea, the Barents Sea, the Kara Sea, and the areas between the islands of the Canadian Archipelago show complicated conditions owing to dominating local influences and inflow of different water masses. Atlantic water, for example, flows into the Barents Sea both to the north of Norway, where it branches out in different directions, and to the north of Spitsbergen, where part of the Atlantic water turns south between Northeastland and Franz Josef Land. The different basins in the Barents Sea therefore contain waters of slightly differing characteristics depending upon their origin.

The currents of the Arctic Mediterranean Sea have been repeatedly mentioned in the preceding discussion of the water masses. Later observations have not materially changed the conception of the surface currents of the Norwegian Sea which Helland-Hansen and Nansen presented in 1909. Their picture is reproduced in fig. 180, according to which the two main currents in that region are the Norwegian Current, representing a continuation of the North Atlantic Current, and the East Greenland Current.

The Norwegian Current, a part of the Gulf Stream system (p. 673), is flanked on the left-hand side by a series of whirls, some of which are probably stationary and related to the bottom topography, whereas others may be traveling eddies. Off northern Norway the Norwegian Current branches, one branch continuing into Barents Sea and another turning north toward Spitsbergen and bending around the northwest of the Spitsbergen islands. The waters of this current have a high salinity and a high temperature, the maximum salinity decreasing from about 35.3 ‰ north of Scotland to about 35.0 ‰ off Spitsbergen and the subsurface temperature decreasing from about 8° to less than 4° in the same distance.

The East Greenland Current flows on or directly off the East Greenland shelf, carrying water of low salinity and low temperature. The greater part of the East Greenland Current continues through Denmark Strait between Iceland and Greenland, but one branch, the East Iceland Arctic Current, turns to the east and forms a portion of the counterclock-wise circulation in the southern part of the Norwegian Sea.

No velocities are entered on the figure, but within the Norwegian Current velocities up to 30 cm/sec, or about 0.5 knot, occur and within the East Greenland Current, where the water moves fastest directly off the shelf, velocities of 25 to 35 cm/sec are encountered. These values are based partly upon computation from the distribution of density and partly upon direct current measurements (Helland-Hansen and Nansen 1909, Jackhelln, 1936).

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In the North Sea a counterclockwise circulation is also present, as has been demonstrated not only by the distribution of temperature and salinity but also by the results of large-scale experiments with drift bottles. The details of the currents are much more complicated than can be shown in the figure, and several smaller but permanent eddies appear to be present. In the straits between Sweden and Denmark the surface current is directed in general from the Baltic to the North Sea, but along the bottom the water flows into the Baltic. In the Baltic and the adjacent gulfs the currents are so much governed by local wind conditions that no generalization is possible.

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The outflowing water from the Baltic continues as a low-salinity current along the south and southwest coast of Norway, as shown in fig. 180. This Norwegian coastal current carries warm water in summer and cold water in winter, the development of the current being greatly influenced by the prevailing wind. A striking example of the advance of cold water masses along the coast in the winter of 1937 is discussed by Eggvin (1940), who traced the movements of a cold-water front by surface thermograph records and oceanographic observations at fixed localities.

In the North Polar Sea the surface currents are also greatly influenced by local winds. An independent current appears to be present only to the north of Spitsbergen and to the northeast of Greenland, where the surface waters flow south to feed the East Greenland Arctic Current. The greater part of the Atlantic water which reaches Spitsbergen as the northern branch of the Norwegian Current submerges below the Arctic surface water and spreads as an intermediate layer over large parts of the Polar Sea.

In the Barents Sea a counterclockwise circulation prevails with relatively warm water of Atlantic origin on the southern side and Arctic water on the northern, and with numerous eddies in the central portion. In the other marginal areas of the Polar Sea—the Kara Sea, the Laptev Sea, the North Siberian Sea, and the Chukotsk Sea—the currents are mainly determined by local winds and, in summer, by the discharge of large quantities of fresh water from the Siberian rivers or the Yukon River, but details of the currents are little known.

The Norwegian Current and its branches are subject to variations which are related to other phenomena. Helland-Hansen and Nansen (1909, 1920) have shown that the surface temperature of the Atlantic water off the Norwegian coast fluctuates considerably from one year to another and that high summer temperatures of the surface water are mostly followed by high air temperatures during the subsequent winter and spring. Besides such minor fluctuations major changes take place that lead to altered conditions over a number of years (Helland-Hansen, 1934). In 1901–1905 and in 1925 and 1927 the maximum salinity of the Atlantic water off western Norway at depths greater than 50 m was between 35.30 ‰ and 35.35 ‰, but in 1929 a maximum salinity of 35.45 ‰ was observed, and in 1928 and 1930 the highest values were 35.43 ‰ and 35.40 ‰ respectively. The temperatures in 1929 were higher than on any previous occasion and, in agreement with this observation, the air temperature in Norway from November, 1929, to April, 1930, was higher, on an average, than in any winter since 1900.

Helland-Hansen found that in 1929 the Atlantic water flowing into the Norwegian Sea was not only warmer and of higher salinity but that the volume was also greater than average. He points out that such an increased inflow must have had far-reaching consequences. Owing to

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the time needed for the water to reach the Barents Sea the effect of the large amount of warm water should appear in the Barents Sea two years after the water was observed off southwestern Norway, and the inflow of warmer water should lead to an increase of the ice-free areas in spring. In agreement with this reasoning, in May, 1929, the ice-free areas in the Barents Sea east of 20°E comprised 330 k m²; in May, 1931, they extended over 710 k m². An effect on conditions to the northwest of Spitsbergen should also appear one or two years later. Continuous observations are not available, but Mosby (1938) points out that, whereas in the years 1910, 1912, 1922, and 1923 the maximum salinity of the Atlantic water varied between 34.95 ‰ and 35.06 ‰, it reached a value of 35.14 ‰ in 1931. In earlier years the maximum temperature varied between 2.57° and 4.48°, but in 1931 it reached 5.04°. The inflow of the warmer water led to the return into Spitsbergen waters of the cod, which had not been caught there in commercial quantities during the preceding fifty or sixty years.

The causes of the fluctuations which have been described are not known, nor is it known whether these fluctuations are periodic in character, and the same statements apply to fluctuations which occur in other regions of well-defined currents.

The large water masses of the Arctic Mediterranean have a high oxygen content, but within the many minor bodies of water which are in restricted communication with the waters of the Arctic Mediterranean the oxygen content varies within wide limits. In general, the Atlantic water of the Norwegian Current has a high oxygen content, but off the Norwegian coast minimum values somewhat below 5 ml/L have frequently been observed. The deep water of the Norwegian Sea contains large quantities of oxygen, in agreement with the concept that this water is rapidly renewed by vertical convection currents which in late winter extend from the surface to the bottom. At a temperature of -1° values up to 7.2 ml/L have been observed, corresponding to a saturation of 88 per cent. Both the Atlantic water that flows into the Polar Sea to the north of Spitsbergen and the deep water that enters over the Spitsbergen-Greenland Ridge is rich in oxygen, as is evident from the observations by Mosby (1938) and Sverdrup (1933). In the deep water of the Polar Sea to the north of Spitsbergen the highest observed value was 6.73 ml/L at a depth of 3,000 m. On the other hand, in coastal areas where the bottom topography or other local conditions are important, very low oxygen values may be found. On the North Siberian shelf, where occasional intrusion of high-salinity water takes place, low oxygen values occur, probably owing to the consumption of oxygen by the organisms on the bottom of the shallow shelf or to decomposition of organic matter which has accumulated on the bottom. Values as low as 1.56 ml/L were observed at a distance of 5 m from the bottom, where the salinity was

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about 32.00 ‰, whereas at 15 m from the bottom, where the salinity was 29.00 ‰, the oxygen content was nearly 8 ml/L (Sverdrup, 1929). In numerous basins in Norwegian fjords the water contains no oxygen but does contain large amounts of hydrogen sulphide (Ström, 1936).

The Labrador Sea and Baffin Bay. The waters of the Labrador Sea are at all depths in free communication with those of the Atlantic Ocean, wherefore the Labrador Sea should be considered not an adjacent sea but a part of the Atlantic. The waters of Baffin Bay, on the other hand, are partly separated from those of the Labrador Sea by the submarine ridge across Davis Strait; but, because of the close relationship between the waters of Baffin Bay and Labrador Sea, and because of the strong influence of similar local factors in both regions, it appears desirable to treat them jointly.

The character of the water masses in the Labrador Sea and Baffin Bay is illustrated in fig. 181, based upon diagrams prepared by Smith et al (1937). The distributions of temperature and salinity are shown in a cross section from South Wolf Island off the coast of Labrador to Cape Farewell on Greenland and in a longitudinal section between latitudes 78°N and 55°N. The locations of the sections are indicated on the small inset map.

The cross section from Labrador to Greenland shows a distribution of temperature and salinity having considerable similarity to the distribution found in the Norwegian Sea. On the continental shelf off Labrador (to the left in the figure) is found water of a temperature which is mostly below 0°C and of salinity less than 34.00 ‰, the lowest values being less than 30.00 ‰. This water, which flows south, corresponds in its character to the waters of the East Greenland Current. The greater part of the section shows water of a uniform temperature which, except for a surface layer showing the effect of seasonal heating, is between 4° and 1.7°, by far the greater part of the water having a temperature between 3.5° and 3.0°. The salinity lies between 34.86 ‰ and 34.94 ‰, the

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highest salinity being found in the deep water at some distance from the bottom. At the surface the salinity is about 34.60 ‰, but the seasonal variations are great and higher values may be found in other seasons of the year. The uniform deep and bottom water of the Labrador Sea corresponds to the uniform deep water of the Norwegian Sea, but has a higher temperature, 3° as against -1°, and in part it has a somewhat higher salinity. Two different water masses are present off the coast of Greenland. Beyond the slope is a body of Atlantic water of salinity higher than 34.95 ‰, with maximum values above 35.00 ‰ and with temperatures higher than 4° and up to 6°. This body of water corresponds to the Atlantic water within the Norwegian Current. Close to the Greenland coast is water of a salinity between 34.00 ‰ and 31.00 ‰ and of a temperature, in summer, about 2°. This represents water of the East Greenland Current, which bends north around Cape Farewell and is mixed with the Atlantic water; continuing the comparison with the Norwegian Sea, it corresponds to the Norwegian coastal water.

The longitudinal section which has been constructed along the center line of Baffin Bay and Labrador Sea demonstrates, in the first place, the uniform character of the waters of the central part of the Labrador Sea. In the second place, the section shows that the water in the Baffin basin represents a mixture of Labrador Sea water and surface water which has been greatly diluted by excess precipitation. The low temperatures near the bottom indicate an admixture of surface water the salinity of which has been increased in winter by the freezing of ice.

The processes of cooling and of mixing which go on in the Labrador Sea are comparable to those in the Norwegian Sea. The large body of relatively uniform water in the central part of the Labrador Sea is formed by mixing of the two essentially different types of water, the Atlantic water and the Arctic water. In certain areas the mixed water at the surface will have a salinity in the neighborhood of 34.90 ‰ and, when cooled in winter to a temperature of 3° or lower, will sink and renew the deep and bottom water. That is, formation of deep and bottom water takes place in winter by vertical convection currents in the manner which was first described by Nansen when discussing the formation of the bottom water in the Norwegian Sea. The fact that the deep and bottom water has an oxygen content between 6.0 ml/L and 6.5 ml/L strongly supports the concept that this water is formed by vertical convection currents and is being renewed every winter.

In the Baffin basin entirely different conditions are encountered. The temperature and salinity of the bottom water in this basin indicate, as already stated, that here the bottom water represents a mixture of Labrador Sea deep water and surface water of the Baffin Bay region, the salinity of which has been increased sufficiently by freezing to cause the water to sink. It is probable that this sinking of cold surface water is

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a slow and intermittent process, because the oxygen content of the bottom water is as low as 3.5 ml/L. It is of interest to observe that the processes by means of which the bottom water of the Baffin basin is renewed are similar to those that take place in the Weddell Sea area, except that there sinking of water which is cooled on the continental shelf takes place more frequently, as is demonstrated by the higher oxygen content of the bottom water. A T-S curve from the central portion of the Baffin basin has a form similar to that of a T-S curve from the Weddell Sea, as should be expected if processes of similar character operate.

The surface currents of the Labrador Sea, which have been briefly mentioned, are shown in fig. 182, prepared by means of the charts of Smith et al (1937) and of Kiilerich (1939). These representations are the results of calculations based on the distribution of density. The outstanding features are the West Greenland Current that flows north along the west coast of Greenland and the Labrador Current that flows south off the coast of Labrador. Part of the West Greenland Current turns around when approaching Davis Strait and joins the Labrador Current, whereas part continues into Baffin Bay, where it rapidly loses its character as a warm current. Along the west side of Baffin Bay the Arctic waters flow south, having been partly reinforced by currents carrying Arctic water through the sounds between the islands to the west of Greenland. The central areas of the Labrador Sea and Baffin Bay both appear as areas in which numerous eddies occur, but nothing is known as to the permanency of the details shown in the picture. The similarity between the Labrador Sea and the Norwegian Sea is again striking, but it should be emphasized that, whereas the Norwegian Sea is in communication with the Atlantic Ocean in the upper 600 m only, so that the deep water is shut off, the Labrador Sea is in communication with the Atlantic at all depths and the deep water can freely flow to the south.

According to Smith et al (1937) the inflow in the Labrador Sea amounts to 7.5 million m³/sec and the outflow along the coast of Labrador amounts

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to 5.6 million m³/sec, both values referring to flow above a depth of 1500 m. From these figures Smith et al conclude that approximately 1.9 million m³/sec sink and flow out from the Labrador Sea as deep water. This conclusion can be further examined by means of the available cross sections from Labrador to Greenland through which the net flow must be zero. The Godthaab section (Kiilerich, 1939) indicates, in August, 1928, an inflow of deep water amounting to about 0.5 million m³/sec, whereas the General Greene section in August, 1935, indicates an outflow of deep water of 3.6 million m³/sec, the outflow taking place below a depth of 1000 m. The discrepancy between the results from the two sections shows perhaps that the outflow of deep water is an intermittent process. If this is true the greater emphasis should be given to the above values calculated by Smith et al, which are based upon observations from a number of years and according to which the average outflow of deep water from the Labrador Sea is about 2 million m³/sec. This flow to the south has important bearing on the entire deep-sea circulation of the Atlantic Ocean, as will be shown later on.

Ice and Icebergs in the Arctic. The Polar Sea with its adjacent seas, the western portion of the Norwegian Sea, Baffin Bay, and the western portion of the Labrador Sea are covered by sea ice during the greater part of the year, whereas, under the influence of the warm Atlantic water, the west coast of Norway is always ice-free except for occasional freezing over of the inner parts of fjords.

The Arctic pack ice that covers most areas is more broken and piled up than the Antarctic pack ice (p. 623). Large, flat ice floes are rare, but fields of hummocked ice and pressure ridges rising up to five or six meters above the general level of the ice are frequently found. The broken-up and rugged appearance of the Arctic pack ice is ascribed to the action of the wind in conjunction with the restricted freedom of motion of the ice owing to land barriers on all sides.

The wind drift of the ice has been discussed by Nansen (1902) and Sverdrup (1928), who found that the direction of the drift deviated, on an average, 28° and 33°, respectively, from the direction of the wind, instead of 45° as required by Ekman's theory of wind currents. The discrepancy is due to the resistance against motion offered by the ice itself (Sverdrup, 1928, Rossby and Montgomery, 1935) and because this resistance is greatest at the end of the winter when the ice is most closely packed, the deviation of the ice drift from the wind direction is smallest at that time of the year. Similarly the wind factor, that is, the ratio between the velocity of the ice drift and the wind velocity, is smaller at the end of the winter than in summer, varying between 1.4 × 10⁻² in April to 2.4 × 10⁻² in September.

Under the influence of variable winds the ice is, in all seasons of the year, torn apart in some localities, where lanes of open water are formed,

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and is packed together in others. In winter the lanes are rapidly covered by young ice, which in a week or less may reach a thickness of 50 cm; but from the end of June to the middle of August no freezing takes place, because over the entire Polar Sea the air temperature remains at zero degrees or a little above zero. Where lanes are formed, the ice always breaks along a jagged line, and when the ice fields move apart they also are displaced laterally. In summer such lanes remain ice-free. When they close, as the ice is packed together, the two sides of the lane do not fit; corners meet corners, and between the corners openings of different shapes remain, many of them several hundred meters long. In summer the Arctic ice fields are therefore not continuous, but are honeycombed to such an extent that from no point can one advance as much as 10 km in any direction without striking a large opening in the ice. This characteristic makes possible the use of a submarine for exploration of the Polar Sea, as advocated by Sir Hubert Wilkins.

In winter the average thickness of the Arctic pack ice is probably 3 to 4 m and in summer 2 to 3 m, but under pressure ridges and great hummocks the thickness is greater. Hummocked ice is often found stranded where the depth is 8 to 9 m, and in exceptional localities where strong tidal currents prevail, masses of piled-up ice have been found stranded where the depth was 20 m.

The ice passes through a regular annual cycle. In summer melting takes place during 2 to 3 months, and on an average the upper 1 m of the ice melts. In winter ice forms on the underside of the floes, but the thicker the ice the slower the freezing. The average thickness of the ice depends mainly upon the rapidity of melting in summer and freezing in winter, and is therefore determined by climatic factors. Near shore, river water and warm offshore winds facilitate melting and the development of navigable lanes of open water along the coasts. In recent years the U.S.S.R. has been able to take advantage of such lanes along the north coast of Siberia for establishing shipping connections with the large Siberian rivers. Aerial surveys of ice conditions have preceded the operations and ice breakers have been used where necessary.

Within the greater part of the Polar Sea the ice is moving mainly under the action of the winds, but it is also carried slowly by currents toward the opening between Spitsbergen and Greenland. The speed of the current increases when approaching the opening, and great masses of ice are carried swiftly south by the East Greenland Current. Since 1894 detailed information as to the extent of the sea ice in the Norwegian Sea and in the Barents Sea has been annually compiled and published by the Danish Meteorological Institute.

The icebergs in the Arctic originate from glaciers, particularly on Northern Land, Franz Josef Land, Spitsbergen, and Greenland. No icebergs are encountered in the Polar Sea except near Northern Land,

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Franz Josef Land, and Spitsbergen because the Greenland glaciers do not terminate in the Polar Sea. The glaciers on the first three islands are small and produce only small icebergs. By far the greater number, and all large icebergs, originate from the Greenland glaciers and are carried south by the East Greenland and Labrador Currents. These icebergs are generally of irregular shape because the Greenland glaciers do not form thick shelf ice comparable to that of the Antarctic, but terminate in fjords where piece after piece breaks off as the glacier advances. Some fjords are closed by shallow sills on which the bergs strand.

Many of the icebergs carried south by the East Greenland Current disintegrate before they reach Cape Farewell, the southern cape of Greenland, but others are carried around the south end of Greenland and continue to the north along the west coast. They are joined by other icebergs from the West Greenland glaciers and together with these are finally carried south by the Labrador Current. The icebergs reach farthest south off the Grand Banks of Newfoundland in the months from March to June or July, when they represent a serious menace to shipping. After the Titanic disaster in 1911, the International Ice Patrol was established in order to safeguard the shipping by reporting icebergs and predicting their probable course. The ice patrol is conducted by the United States Coast Guard, the publications of which contain a large amount of information as to the number, distribution, and drift of icebergs in the region of the Grand Banks. Prediction of the ice drift has been based successfully on currents computed from the distribution of density as observed on cruises during which several lines of oceanographic stations have been occupied in about two weeks. This work of the Coast Guard is the outstanding example of practical application of the methods for computing ocean currents (p. 453).