Weather and Climate
Douglas R. Powell and Harold E. Klieforth
The highest range of the many mountain ranges that are arranged en echelon in the Great Basin between the Sierra Nevada and the Wasatch Mountains is the White Mountains, situated along the California-Nevada border about 225 mi (362 km) east of the Pacific Coast. Climatically, this location is transitional between the moderating maritime influence of the Pacific Ocean and the more extreme continental influence of interior North America. The very large variation in elevation within the range, from 4,000–5,000 ft (1,200–1,500 m) at the base to 14,246 ft (4,343 m) at the summit, results in rapid and significant changes in temperature and precipitation within short horizontal distances. Any air mass reaching the White Mountains must pass over an assemblage of other ranges that vary in width and altitude. By far the most important of these mountain barriers is the equally high Sierra Nevada, lying immediately to the west, which, by inducing air to rise, clouds to form, and precipitation to fall, intercepts some of the moisture from Pacific storms in the winter half of the year. To the north, east, and south are a series of lower ranges that have lesser climatic influence. Topographically, the least impeded avenue of approach for an air mass is from the southeast, but significant movement of air from this direction is uncommon. When it does occur, normally in July or August, the result may be spectacular thunderstorms with high precipitation intensities.
Atmospheric Circulation Patterns
Unlike other features of the physical environment, the gases that constitute the atmosphere are invisible, so that it is necessary to use indirect means to describe graphically and to map continuously changing atmospheric conditions. Atmospheric motions are complex, but when studied they are found to follow patterns. To understand the weather regimes of the White-Inyo mountains, it is helpful to recognize three principal scales of motion, each roughly an order of magnitude greater than the next.
The first and largest of these is the synoptic scale , so called because it is analyzed from numerous soundings, measurements, and observations made at the same time at hundreds of locations around the world. This is the scale of the familiar weather maps seen in daily newspapers and on television. These surface and upper-air charts cover horizontal distances of 100 to 2,000 mi (160 to 3,200 km) and depict such features as low- and high-pressure areas, cyclones, and anticyclones, air masses, weather fronts, and regions of precipitation. Generally, these features progress in a predictable manner,
and from them it is possible to produce, by computerized prediction models, future patterns of airflow, moisture, temperature, clouds, and precipitation. These are used to forecast, with varying degrees of certainty, weather conditions for a particular region for a few days following receipt of the climatic data.
The second scale of atmospheric motion, the mesoscale , describes airflow patterns over distances of, say, 1 to 100 mi (or 1.6 to 150 km), a range that includes many of the spectacular cloud formations and weather conditions experienced in and near the mountains of eastern California (Plates 1.1–1.8). These phenomena are closely related to the synoptic flow patterns but are controlled and shaped by major terrain features such as the Sierra Nevada and the White-Inyo Range.
The third scale, the toposcale , applies to weather and climatic conditions within a distance of, say, 1 mi (1.6 km) that vary in relation to prominent features shown on local (15-minute or 7.5-minute series) topographic maps. Thus, under differing synoptic and mesoscale conditions different air temperatures, wind velocities, and snow accumulations are measured, whether the location is on a mountain ridge or in a canyon, on a windward or leeward slope, or in a broad valley.
In the sections that follow we discuss the principal patterns that affect the White-Inyo Range, its inhabitants, and its visitors in each of these different scales. We begin with a survey of synoptic-scale airflow and its relation to seasonal weather.
Figure 1.1 shows the principal airflow patterns and air-mass types or source regions that determine regional weather in different seasons of the year. The arrows indicate the directions of air movement near and above the crest of the major mountain ranges, at levels between 10,000 and 20,000 ft (3 to 6 km) above sea level. The open circles are locations from which twice-daily (near 4 A.M. and 4 P.M. PST) rawinsonde balloon ascents are made to obtain data on air temperature, humidity, and wind velocity, from which upper-air (e.g., 500 mb) weather maps are plotted and analyzed. The black circle indicates the location of the Bishop Airport, the nearest (3 mi, or 5 km) National Weather Service station to the White-Inyo Range.
The air that flows across California at any time of year is most likely to have passed over some part of the Pacific Ocean. In summer the Pacific Anticyclone (a large, slow-moving clockwise whirl of air) lies just west of California, bringing an onshore flow of cool marine air, stratus clouds, and fog to the coast and mostly clear, dry air to the Sierra and White-Inyo Range. During much of the summer the Great Basin Anticyclone develops over the warm plateau region of Nevada and Utah. When this whirl expands and shifts westward, a flow of moist maritime tropical air from the Gulf of California or the Gulf of Mexico may persist for a few days before the normally dry Pacific flow reasserts itself. Thus, during the summer season the mountainous terrain of eastern California and western Nevada is contested for by two air masses, with that from the northern or central Pacific usually prevailing.
During fall, winter, and spring a series of traveling upper-air troughs (cyclonic bends with counterclockwise flow) and ridges (anticyclonic bends with clockwise flow) cross California and Nevada. Ridges and anticyclones usually bring subsiding air, few clouds, and "fair" weather, whereas the troughs and associated cyclones, low-pressure centers, and fronts bring much cloudiness and widespread precipitation. Fronts are boundaries between converging air masses from different source regions. The primary air masses affecting California are cold maritime polar air from the Gulf of Alaska and warmer, moist maritime subtropical air from lower latitudes. Occasionally there are invasions of cold continental polar air from northern Canada or the Rocky Mountains.
Cloud formation and precipitation result primarily from ascending motion in moist air. When air rises, it expands and cools. This causes its invisible water vapor to
condense, forming small cloud droplets. As the ascending motion continues, the clouds thicken and the droplets grow larger to form raindrops. If it is sufficiently below freezing (32°F, or 0°C), snow crystals form, which, when heavy enough, fall from the clouds and reach the ground as rain or snow. The precipitation that falls on the mountains each year is a result of four principal mechanisms of upward-moving air: general ascent produced by widespread horizontal convergence in cyclonic flow, more intense lifting in frontal zones, strong "orographic," or terrain-induced, lifting over the windward slopes of mountains, and thermal convective instability triggered by the ascending motion, which causes the cloud to billow upward and precipitate with greater intensity. Although heavy precipitation (40 to 80 in or 100 to 200 cm of water annually) falls on the upper western slope of the Sierra, the region immediately leeward, including the Owens Valley, the White-Inyo Range, and much of western Nevada, is in the so-called rain shadow of the Sierra and receives much less precipitation. The drying of the air results from both the loss of moisture due to precipitation on the windward slope and the adiabatic warming induced by descent over the leeward slope.
Cold fronts, moving in such a way that cold air replaces warmer air at the surface, usually approach from the northwest or the west. Before the passage of the front and its lagging upper-air trough, the airflow across the mountains is from the west or southwest, while surface winds in the valleys and basins are mild and from the south. After frontal passage the upper flow becomes northwesterly to northerly, and cold northerly to northeasterly winds sweep across the mountain ridges and along the valleys. The Sierra Nevada has a profound effect on most fronts, causing them to stall west of the crest while their northern sectors move more rapidly across Oregon and then southward over northern Nevada. Thus, many fronts converge on the Sierra Nevada and White-Inyo Range in nutcracker fashion, with the cold air reaching the leeward valleys and basins from the north before the cold air from the west can surmount the High Sierra.
When a surface low-pressure center forms in western Nevada accompanied by an upper trough that deepens excessively to form a cyclone over the region — a weather pattern known locally as a "Tonopah low" — a northeasterly to southeasterly flow often brings continental polar air or recycled maritime polar air, low clouds, and snowfall to the White-Inyo Range. When such storms involve moist Pacific air, they usually bring heavy snowfall to the region; some of the biggest snowstorms recorded in the White Mountains have occurred in such circulation patterns. Precipitation from closed cyclones over the region is most frequent in spring, resulting in a spring (April or May) precipitation maximum in much of the Great Basin, in contrast to the pronounced winter maximum in the Sierra Nevada.
On infrequent occasions, usually several years apart (e.g., January 1937, January 1949, December 1972, February 1989, and December 1990), a long northerly fetch of air may bring an invasion of true Arctic air from interior Alaska or the Yukon. These episodes bring record cold temperatures to the White-Inyo Range and adjacent valleys; at such times minimum temperatures may dip to -25°F (-31°C) or below.
Most winters include one or two episodes of "warm storms," periods of a few to several days in which very moist tropical air reaches California from the vicinity of Hawaii. In these events the freezing level may be above 10,000 ft (3,000 m), and the heavy rainfall may result in widespread flooding in much of California. It is during such storms that heavy rime icing may form on trees, structures, and power lines on high mountain ridges. This is caused by the combined effect of strong winds and supercooled clouds (composed of water droplets at air temperatures below freezing). The cloud droplets freeze on contact, building great formations of ice that grow into the direction of the wind.
Conversely, "cold storms" bring snow to low elevations, including the floor of Owen Valley and desert areas to the south and east of the White-Inyo Range. Major westerly storms that last for two or three days bring heavy accumulations of snow — a foot (30 cm) or more in the valleys, and two or three times as much at the highest elevations. Very cold storms from the northwest contain less water vapor, are of shorter duration, and usually bring only a few inches (several centimeters) of snow.
Mountain Lee Waves
When a cold front approaches California from the northwest and the westerly airflow increases in speed over the Sierra crest, spectacular "stationary" clouds are usually seen over the leeward valleys. These are manifestations of a mountain lee wave, as it is known (Fig. 1.2). If the ridges and troughs of the horizontal airflow pattern are likened to the bends or meanders of a stream, the lee wave phenomena are analogous to the falls and ripples. Figure 1.3 shows a typical pattern of airflow and cloud forms in a strong lee wave. Air flowing over the Sierra Nevada plunges downward, then upward, and then downward again in a series of crests and troughs. The wavelength depends on the airflow characteristics, mainly the variation of air temperature with height (lapse rate) and the increase of wind speed with height (wind shear). The amplitude is greatest in strong waves and in cases where the vertical flow pattern is in resonance with the terrain, as, for example, when the second wave crest lies over the next mountain range downwind, such as the White-Inyo mountains, east of Owens Valley.
Updrafts and downdrafts in a strong lee wave often have speeds of 2,000 ft (600 m) per minute, sometimes exceeding 4,000 ft (1,200 m) per minute. Where the air descends, it warms, and the relative humidity decreases. The warm, dry winds, which may reach speeds of 60 mph (30 m/s) or greater at the surface, are known as foehn winds. The stratocumulus cloud deck over the Sierra Nevada is called a foehn-wall or cap cloud, and its downslope extension is known as a cloudfall. After it evaporates, the invisible moisture is cooled again in the ascending current and forms the turbulent cumuliform roll cloud. Looking at the roll cloud, an observer has the impression that it is rotating, but this is an illusion caused by the wind shear. However, below the roll cloud there is commonly a true rotor circulation, which brings easterly winds at
the surface. On rare occasions a very strong lee wave will be in resonance with the terrain so that the first wave crest lies over the White-Inyo Range.
Above the roll cloud, there may be one layer or several decks of smooth, lens-shaped altocumulus clouds which appear stationary but through which the air is passing at 50–100 mph (22–45 m/s); the cloud droplets form at the windward edge and evaporate at the leeward edge. Soaring pilots make use of a mountain lee wave by flying into the wind and ascending in the updraft zone. In Figure 1.3 the dotted line shows a typical path of a sailplane: a line of flight under the roll cloud during airplane tow, release point at "x," and then a line of flight upward. Several flights above altitudes of 40,000 ft (12,000 m) have been made in lee waves by pilots equipped with oxygen to survive the low pressure (200 to 150 mb) and with warm clothing to withstand the cold temperatures (-94°F or -70°C).
Wave clouds may appear in any month of the year, but they are most often seen in late winter and in spring. They usually reach their maximum development in
midafternoon and are most beautiful at sunset, when the highest clouds, at 30,000–40,000 ft (9,000–12,000 m), remain colorful long after the sunlight has left the leeward valley.
A local wind of another sort may sometimes be observed by motorists in the winter season in the Owens Valley near Olancha. During strong lee wave conditions or the passage of a cold front, a great horizontal cyclonic eddy may develop about Owens (dry) Lake; there will be a northerly (i.e., from the north) wind along the route of Highway 395 but, as evidenced by blowing dust, a southerly wind on the east side toward Keeler. When the surface flow is southerly in Owens Valley, the dust from Owens Lake may be carried far north of Bishop, lowering visibility so that the mountains are nearly obscured, and the alkali dust is sometimes tasted by pilots flying at 12,000 ft (3,600 m)! Similar phenomena occur in other arid basins of California and Nevada.
At any time from early May to early October, there may be an incursion of tropical air from the south; then thunderstorms are possible. The intense heating of the arid Southwest during the summer months creates the upper-air anticyclone and surface low-pressure area that provide the circulation necessary for the northward flow of tropical air. In eastern California and Nevada, this summer monsoon is best developed during the period from early July to late August.
At first, the moisture enters the area at the high levels, and a thundery spell is commonly heralded by the appearance of rather exotic cirrus clouds from the south quadrant. Within a day or two, if the flow persists, the air at middle and low levels is also moist, and daily thundershowers can occur. These are strongly diurnal in their development; that is, they develop as a result of the daily heating of the mountain slopes by the sun, and they decline after sunset.
The appearance of patchy, turreted altocumulus clouds at sunrise is a good indication of possible thunderstorms later in the day. Heating of the rocky mountain slopes causes the air to rise toward the crests, and soon cumulus clouds form above these upslope currents. The clouds continue to rise upward, becoming what the weather observer calls towering cumulus. Near midday their tops develop a fibrous appearance indicative of ice crystal formation, and they are said to be glaciated. Soon they develop anvil-shaped cirrus tops with streamers of ice clouds stretching downwind at those levels (30,000–40,000 ft, or 9,000–12,000 m). At this time, lightning flashes from cloud to mountain and heavy local showers of rain and, commonly, graupel , or pea-sized hail, fall on the crests and ridges of the ranges. By this time hikers and climbers should have taken shelter.
Later, as downdrafts of cool air predominate, the thunderstorm ceases, and as the sky brightens to the west, the clouds begin to thicken over the leeward valley, where the day-long heating has created rising thermal currents. Often at 5:00 P.M. or 6:00 P.M. PST a brief thundershower is experienced in valley locations. As the shower moves eastward and the lowering sun in the west shines on the dark cloud and rain shafts, a brilliant rainbow is visible. Finally, when the sky has cleared and stars have appeared, lightning might continue to flash in the east if a nocturnal storm continues over central Nevada.
During some summers there are numerous thunderstorms, as in 1955, 1956, 1967, 1976, 1983, and 1984; in other years there are very few. It is difficult to predict exactly where the storms will occur, as this depends on subtle differences in wind velocity, amount of moisture, rate of growth, and topography. In the morning hours, the eastern slopes of the mountains are heated, the warm currents rise, and an easterly upslope wind forms in the valleys. If the upper synoptic flow has an easterly component, the clouds will develop even more rapidly. In the afternoon, the western slopes of the mountains are heated more effectively by the sun, and, especially if aided by a westerly breeze, the cloud development intensifies there. Once the cloud has formed, latent heat is released, which increases the cloud's buoyancy and causes it to rise more vigorously. Because these storms are so localized, commonly affecting a single canyon, intense cloudbursts may cause flash floods. These commonly occur when cloud bases are below the mountain tops. Such events often go unobserved in the Inyo Mountains and remote Nevada ranges and are discovered some days or weeks later. The damage is usually greater there, though, because roads commonly follow the canyons and washes.
As a general rule, it does not rain on summer nights in the mountains, but there are exceptions. Occasionally, the remnants of tropical storms called "easterly waves"
are carried northward along the Sierra and over much of Nevada. The cloudiness is general, and precipitation may be widespread, continuing at night and commonly accompanied by low clouds and fog, lightning, and thundershowers. Such episodes occurred in July 1956 and in August 1965.
The hiker's pocket altimeter and the altimeter in an aircraft are merely barometers measuring atmospheric pressure and indicating the equivalent height according to average weather conditions. Pressure, which always decreases with altitude, is usually measured in millibars, as shown at the right side of Table 1.1, or in inches (or millimeters) of mercury. Some equivalent pressures and heights for average atmospheric conditions in the California-Nevada region are listed in Table 1.1. The physiological effects of increasing altitude and decreasing pressure are mainly caused by the reduced amount of oxygen and the greater effort one has to exert to get enough oxygen into the lungs. Most hikers have to become acclimated for a day or two before they can be comfortable above 10,000 ft.
At most times and places, air temperature also decreases with height because the atmosphere is mainly heated from below. The atmosphere does not absorb much of the direct radiation from the sun, but the earth's surface does and reradiates the energy at a longer wavelength, which the atmosphere absorbs mainly through two of its variable gases: water vapor and carbon dioxide. The average lapse rate (generally, the decrease of temperature with altitude) is approximately 3.6°F per 1,000 ft (6.5°C per km). At midday in summer with strong thermal activity, the lapse rate on the slopes approaches the adiabatic value of about 5.4°F per 1,000 ft (10°C per km) of ascent. Thus, on days when the temperature is 90°F (32°C), for example, in the Owens Valley, it can be 45° to 55°F (7° to 13°C) on Mt. Whitney or White Mountain Peak (both above 14,000 ft or 4,000 m). On clear nights, on the other hand, cooler air with its greater density sinks and collects in the valleys, forming inversions in which the temperature increases with height in the lowest few hundred feet (» 100 m) above the ground.
Topography influences local weather and climate in many ways, examples of which will be noted throughout the following sections of this chapter.
Weather Observations and Climatological Data
Sparseness of permanent population, little human use of the area, and inaccessibility of the terrain have led to a scarcity of reliable weather records. Students and aficionados of mountain weather and climate always bemoan the paucity of weather instruments and observers at high elevations. Figure 1.4 shows the locations and elevations for weather stations in and immediately adjacent to the White Mountains. Table 1.2 gives the average monthly and annual temperatures for these stations, and Table 1.3 lists the corresponding precipitation data. Unfortunately, there are gaps in the records from most of the stations, especially during extreme weather events when such data can be very useful. Most reliable is Bishop (see Fig. 1.4), at 4,108 ft (1,250 m) in Owens Valley to the west of the White-Inyo Range, operated by the National Weather Service since 1947. The other stations are maintained by cooperative agencies, institutions, or individuals. The only continuous records from within the mountains proper are from stations operated by the White Mountain Research Station (WMRS) — White Mountain I in the valley of Crooked Creek, at 10,150 ft (3,095 m), and White Mountain II on the east slope of Mt. Barcroft, at 12,470 ft (3,800 m). Regrettably, maintenance costs and problems have closed White Mountain II during the winter months from January 1980 to the present, and the record from White Mountain I stopped after 1977. Automated weather-recording equipment is now being installed in the White Mountains by the WMRS. Dyer, at 4,975 ft (1,517 m) in Fish Lake Valley, and Deep Springs College, at 5,225 ft (1,593 m) in Deep Springs Valley, are representative of the lowland valleys to the east and southeast of the mountains. Benton, 5,377 ft (1,640 m), and Basalt, 6,358 ft (1,940 m), to the northwest and
north of the mountains have incomplete records and were little used in this climatic analysis. Unless otherwise noted in the tables, the period of record is from 1956 through 1985; 30 years is generally considered by climatologists to be the minimum length of time necessary to establish meaningful averages.
Following is a summary of three important components of weather and climate in and around the White Mountains: temperature, precipitation, and wind. Data from the above-mentioned stations serve as a basis for the discussion. Monthly average temperatures are calculated by averaging the daily maximum readings for the month, averaging the daily minima, adding these two totals, and dividing by 2. In the following discussion, winter includes December, January, and February; spring March, April, and May; summer June, July, and August; and fall September, October, and November. Both authors of this chapter have many years of direct observation
of weather events in the White Mountains and have used personal experience and knowledge to augment interpretation of the formal record, especially in the large portions of the range not covered by the recorded data.
The summer visitor to Owens Valley at midday may well think the valley and the apparently barren, flat-lighted slopes of the White Mountains to the east to be under the full possession of the sun, with a forbidding aspect of heat. The same visitor to elevations above 10,000 ft (3,049 m) would likely comment — perhaps complain — about coolness, or even the cold. In July, the warmest month, average daily temperatures are 70°F (21°C) or higher on all the adjacent valley floors; the maximum daily temperature reached 109°F (43°C) at Bishop in June 1969 and July 1982, 106°F (41°C) in June 1961 at Dyer, and 104°F (40°C) at Deep Springs in June 1964. The July average is 51.7°F (10.9°C) at White Mountain I and 46.7°F (8.2°C) at White Mountain II, with maximum summer temperatures reaching 79°F (26°C) at I in July 1967, and 73°F (23°C) at II in August 1978. The decline in average July temperatures with an increase in altitude is close to the 3.6°F drop per 1,000 ft (6.5°C per km) rise in elevation regarded as the normal lapse rate throughout the world. It is a rarity for any of the summer months at either mountain station not to have one or more readings of 32°F (0°C) or below, and even the adjacent valleys have such readings in June and August. For any location on earth, trees are generally absent if the average of the warmest month is below 50°F (10°C). This critical value occurs between White Mountain I and White Mountain II; the former has trees, and the latter has none.
In January, the coldest month, the average temperature is 37.2°F (2.9°C) at Bishop, with a winter low of -7°F (-21.5°C) in January 1982. Temperatures are somewhat colder in the valleys to the east, with Dyer at 31.4°F (-0.3°C) with a winter low of -21°F (-29.5°C) in January 1962 and 1974, and Deep Springs at 30.6°F (-0.8°C), with a low of -10°F (-23.5°C) in January 1973. Fish Lake and Deep Springs Valley are colder than Owens Valley because they are at higher elevations, they are farther from the maritime influence of the Pacific Ocean, and they are more open to invasions of cold air from the north and east. At White Mountain I January is the coldest month, with an average of 20.6°F (-6.3°C) and a winter low of -25°F (-31.5°C) in March 1968. February is the coldest month at White Mountain II, with an average of 14.8°F (-9.6°C) and a winter low of -35°F (-37°C) in March 1964. In winter there is less of a decrease in monthly average temperatures with a rise in elevation than in summer. Cold air settling in the lower valleys at night from radiational cooling and downslope drainage seems to be the major reason: the temperature difference between the valleys and the slopes or highlands is less pronounced for minimum readings than for daily maxima, especially on clear nights. This phenomenon is characteristic of high mountains everywhere. At White Mountain I, located in a valley, average minima for the three coldest months
(January, February, and March) are 2°F (1°C) warmer than at White Mountain II, which is 2,320 ft (707 m) higher on a slope; average maxima for those same months are 9.5°F (5.3°C) higher. Thus, a typical early winter morning at the lower station would be just about as cold as at the higher station, but midafternoon would be noticeably warmer.
There is also a much greater variation in average monthly temperatures in winter than in other seasons at all stations and elevations. At some time during most winters, there is an invasion of continental polar or arctic air from northern Canada and Alaska, which brings below-normal temperatures. The frequency and duration of these incursions of cold air vary greatly from year to year; on rare occasions they dominate the weather for weeks. The last such major occurrence was in January and February 1949, and in January 1937 before that. In 1937, a minimum of -42°F (-41°C) was recorded in January at Fish Lake Valley; readings close to that may have occurred within the mountains. In general, the White Mountains are protected from common and prolonged cold air invasions by the many mountain ranges to the north and east.
February and March averages are lower than those for January at White Mountain II, and they are only slightly higher than January averages at White Mountain I. This is not the case at the lowland stations in the area or at most inland stations in cold-climate regions elsewhere in the United States, where January is nearly always the coldest month, with definite warming occurring in February and, particularly, March. At the two high-elevation stations March is colder than December, an anomaly for the latitude. A continous snow cover during the winter and spring could partially account for this delay in warming at high elevation, but other regions in the United States with persistent snow cover do not show this effect. It is probably the result of the frequent passage of closed low-pressure systems in late winter and spring over the White Mountain area. An analysis of 500 mb (near 18,000 ft or 5,500 m) weather maps shows that these closed lows commonly contain very cold air, which could affect the highest elevations of the White Mountains. At lower elevations the increased solar radiation from longer days and a higher angle of the sun would offset the influence of the cold air aloft. Thus, there is a pronounced lag in temperature increase in late winter and early spring at high elevation in the White Mountains. Significant warming there usually does not occur until mid-May, when the frequency and intensity of the upper-level closed lows diminish. Major cooling throughout the area generally comes in late September and October, at a slower rate at high elevation than the warming in late spring, with a marked decrease in November at all stations.
Midwinter temperatures at White Mountain I are comparable to those in central Iowa, and at White Mountain II to southern Minnesota or Anchorage, Alaska. Summer temperatures at White Mountain I are similar to those at the northern limit of trees in Alaska and Canada, and at White Mountain II to the treeless Arctic Slope of Alaska. At comparable elevations and latitudes, temperatures in the White Mountains are generally warmer in winter and cooler in summer than in Utah or Colorado, reflecting less maritime influence farther inland. From meager data and personal ex-
perience, there seem to be no significant temperature differences between the White Mountains and the Sierra Nevada at similar elevations, although the Sierra might be expected to have warmer winters and cooler summers because of closer proximity to the Pacific Ocean. It is difficult to compare climates in different mountain ranges because of topographical variations at individual recording sites — north slope, south slope, valley, orientation to prevailing winds, and other factors.
As discussed previously, precipitation in the White Mountain area results primarily from the passage of cyclones with associated fronts during fall, winter, and spring; from closed cyclones in late winter and spring; and from the flow of moist tropical air from the southeast to the southwest quadrant in the summer. Annual amounts vary from 5–6 in (125–150 mm) on the valley floors to 20 in (508 mm) or a little more at the highest elevations. Totals appear to increase right up to the crest of the range. The rate of increase averages about 1.5 to 2.5 in per 1,000 ft (120–205 mm per km) rise. However, this average is difficult to apply to any one portion of the range, and the increase is not linear, being higher at upper elevations. Table 1.3 gives average monthly and annual precipitation amounts for stations within the region.
From west to east in the White Mountain area, there are important differences in the seasonal distribution of precipitation. Bishop, on the west, has the typical regime of most California stations: winter wet and summer dry, with January the wettest month. White Mountain I and White Mountain II have precipitation much more evenly distributed throughout the year. At both mountain stations January is the wettest month, but only by a slight margin. There is no pronounced dry season; June is the driest month, reflecting the gap between the cyclones of winter and spring and the thunderstorms of July and August. Early fall is relatively dry, with a gradual buildup of precipitation to the winter months. Deep Springs, at the southeast edge of the White Mountains, has maximum precipitation amounts in January and February, with a minor peak in July and August, and minimum amounts in June and October. On the east margin of the White Mountains, Dyer has a slight maximum in spring and mid-summer and a minor minimum in December and January. At lower elevations, the western slope of the White Mountains is relatively open to cyclones from the west in winter, partially subject to closed cyclones from the north in the spring, and somewhat protected from thunderstorms in the summer. The eastern slope, in the double rain shadow of the Sierra Nevada and the White Mountains, is protected from winter cyclones but is more open to closed cyclones in spring and thunderstorms in summer. Upper elevations in the White Mountains are relatively open to all three types of storms and show a trimodal maximum of precipitation. Thus, the White Mountain range is truly transitional in seasonal distribution of precipitation between the winter maximum of California and the Sierra Nevada and the more even annual distribution of the eastern Great Basin and Rocky Mountains.
Yearly precipitation totals not only increase with higher elevation in the mountains but very likely are larger in the northern part of the range. There are no station
records to substantiate this assertion, but the experience of many long-time residents of the area and of both writers suggests that the portion of the range from White Mountain Peak north to Boundary Peak receives more precipitation from cyclones and thunderstorms than the region south of the main peak. Occasional measurements made with a standard snow sampler at comparable elevations show greater depth and water content in the snowpack in the northern part of the range. There is also higher streamflow, more extensive former glaciation, and a less xerophytic vegetation north of White Mountain Peak. In the nearby Sierra Nevada, snow survey records show a general decrease in precipitation from north to south, reflecting a lower frequency of passing cyclones. This could also affect the White Mountains. Moreover, the crest of the Sierra Nevada is lower opposite the northern half of the White Mountains than opposite the southern half, and this may allow more moisture to reach across to the northern segment of the White Mountain Range. Still another possible effect of the Sierra Nevada is that, as previously mentioned, fronts may be retarded in crossing the massive barrier of that range, bringing in cooler air from the north and northeast, which may strengthen the fronts and increase precipitation in the northern portion of the White Mountains.
Empirical observation also indicates that the buildup of cumulonimbus clouds in summer thunderstorms is more likely to occur over specific portions of the summit upland than at random. Topographic influence on air moving into the area from characteristic directions is the probable cause. This could add a checkerboard pattern of precipitation distribution independent of more general patterns, such as the increase with elevation and from south to north. Four areas of cloud concentration are noticeable. From south to north, these are Sheep Mountain-Piute (or Paiute) Mountain, the plateau just south of White Mountain Peak, Chiatovich Flats and the area just north of the Cabin Creek-Birch Creek saddle, and the northern portion of Pellisier Flats at the head of Chiatovich Creek. Common features of the four areas are rises in elevation from south to north and broad lateral extent from west to east. Cumulonimbus clouds may form over any part of the range on any summer day, and during extensive storms all or most of the higher elevations may be cloud-covered, but initial formation and greater subsequent development more commonly occur over these four areas.
There are significant departures from normal in amounts of precipitation from month to month and year to year at all elevations in the White Mountains. Most weather stations in the United States use the calendar year in calculating annual amounts. This causes problems in much of California, with its winter-wet, summer-dry regime, and in high-mountain regions, where much of the significant precipitation falls as snow. Thus, the annual snowpack begins in the fall of one calendar year and builds to a maximum in late winter or spring of the next calendar year. Most California stations use a 1 July–30 June precipitation year to avoid this problem. An even better breakdown for the White Mountain Range is to use the water year employed by many hydrologists — 1 October–30 September. This has the advantage of including snow buildup and important July and August precipitation in one annual total, thus giving a more accurate figure of the water available for streamflow and plant growth, much of which occurs from July to September.
The following discussion uses the 1 October–30 September year to show extremes and variation from normal (see Table 1.4). Thus, the year mentioned in the table ends on 30 September and includes the precipitation from October through December of the previous year. Bishop and, by inference, the lower western slope of the range show the largest departures from normal, with the wettest year (17.28 in or 43.9 cm) in 1969 and the driest (1.68 in or 42.5 mm) in 1960. This is a range of 308% to 30% of average. Fish Lake Valley and Deep Springs Valley to the east and southeast show less variation, with both Dyer and Deep Springs ranging from about 190% to 40%. White Mountain I had maximum precipitation in 1967 (26.59 in, or 67.55 cm) and minimum in 1960 (5.57 in, or 14.15 mm), a range of 206% to 45%. At White Mountain II the high total was 33.56 in (85.35 cm) in 1967, and the low was 9.51 in (24.15 cm) in 1960, a range of 187% to 53%. It is to be expected that Bishop, with its low annual average, would have a greater variation from normal than the mountain stations, with higher averages. But Bishop also varies more than the other lowland stations. Bishop and the lower western slope of the range receive most of their rain and snow from winter cyclones, and precipitation totals reflect seasons of frequent or sporadic passage of such storms. The eastern valleys and lower slopes get relatively more moisture from spring and summer storms, and the upper elevations are more open to precipitation in all three seasons. At all stations 1967 and 1969 were very wet, and 1960 was the driest year. Unfortunately, records for the obviously wet years of 1982 and 1983 are incomplete or missing at the mountain stations; both years brought high totals to lowland stations. It is noteworthy that neither White Mountain I nor White Mountain II was very dry in 1976 and 1977, critical drought years in central California. At both stations spring and summer precipitation partially
made up for winter deficiencies. It seems probable that higher elevations in the White Mountains are less subject to either very dry or very wet years than the neighboring Sierra Nevada and Owens Valley; there is less chance that all three types of storms will be common or rare in any one precipitation year.
Extreme monthly totals at Bishop vary from 8.93 in (22.7 cm) in January 1969 to 0.00 for all months of the year. At Dyer, the extremes are 3.44 in (8.75 cm) in August 1983 and 0.00 for all months; at Deep Springs, totals vary from 4.86 in (12.35 cm) in August 1983 to 0.00 for all months. At higher elevation, White Mountain I shows a maximum of 7.53 in (19.1 cm) in December 1966 and a minimum of no rainfall or a trace for all months but January; White Mountain II shows a high at 8.55 in (21.7 cm) for December 1966 and a low of no rainfall or a trace for all months but February. However, summer thunderstorms at high elevation have certainly exceeded these monthly totals. In part of July 1955, one of the authors (D. R. Powell) measured over 11 in (28.0 cm) in a standard rain gauge at Chiatovich Flats, between 10,000 ft (3,050 m) and 11,000 ft (3,350 m), 8.48 in (21.55 cm) of which fell in 2 1/2 hours on July 23. This is the greatest 24-hour total yet recorded in the White Mountains, although it probably has been approached or exceeded during other summer thunderstorms in the range. At the two mountain stations, maximum summer 24-hour totals are about 2 in (50 mm). It is evident that neither station has yet been in the direct path of the most intense thunderstorms. White Mountain II has a winter high 24-hour sum of 4.40 in (11.2 cm) on 6 December 1966, and White Mountain I recorded 3.80 in (9.65 cm) on the same date. Dyer and Deep Springs each show about 2 in (50 mm) for the maximum daily total for any season; Bishop has received more than 3 in (7.6 cm) in each of the three winter months, from Pacific cyclones.
At elevations above 10,000 ft (3,050 m), over 80% of the mean annual precipitation falls as snow. On the valley floors, from 15 to 25% of the average precipitation is snow, with wide fluctuations from year to year in the snow-to-rain ratio. Regardless of the temperature of air masses moving onto land, rainfall is rare from November through April at White Mountain I and practically nonexistent from October through May at White Mountain II. Moreover, snow has been recorded, usually in amounts of less than 6 in (15 cm), in all of the warmest months from June through September at White Mountain II (see Table 1.5 for a summary of snowfall).
Average annual snowfall is low at the base of the mountains, with 9 in (23 cm) at Bishop, 12 in (30 cm) at Dyer, and 15 in (38 cm) at Deep Springs. It builds to 106 in (270 cm) at White Mountain I and 164 in (417 cm) at White Mountain II. As discussed earlier, annual totals are very likely higher at upper elevations north of White Mountain Peak. Maximum seasonal snowfall amounts have been 50–60 in (125–150 cm) at Bishop and Deep Springs, and a little less at Dyer, with low average precipitation in winter. Bishop has had a few seasons with no snowfall; the other lowland stations have had at least 1 in (2.5 cm) or more of snow in all years of
record. At higher elevations measurable snowfall certainly occurs every year. Seasonal totals, 1 October to 30 September, range from 170 in (432 cm) in 1969 to 48.5 in (123 cm) in 1960 at White Mountain I, and from 238 in (605 cm) in 1969 to 83 in (211 cm) in 1960, at White Mountain II. Maximum monthly falls are 76 in (193 cm) at White Mountain I, e.g., in January 1969, and 86 in (218 cm) at White Mountain II in December 1966. Maximum 24-hour totals are 38 in (97 cm) at I and 44 in (36.37 cm) at II, both on 6 December 1966. At White Mountain II, 76 in (193 cm) of snow fell on the two days 5–6 December. Daily accumulations of 10–24 in (25–60 cm) are not infrequent at higher elevation. These are impressive figures for any location in the world, though exceeded in the nearby Sierra Nevada, and winter travelers in the White Mountains should be aware of the difficulty and danger of being out in such intense snowstorms, commonly accompanied by high winds. Where snow has been measured for water equivalent in both stations, or elsewhere at high elevations, a ratio of about 10 in (25 cm) of snow, as it falls, to 1 in (2.5 cm) of water is common. In lieu of actual measurements of melted snow, a ratio of 10-to-1 (snow depth to water equivalent) has been used in the records from all stations. A 10% density is occasionally too high, but less commonly too low, for snowfall at the two mountain stations.
Continuous snow cover at elevations above 10,000 ft (3,050 m) usually begins in late October or mid-November but can begin as early as the end of September and as late as February. Disappearance of snow cover usually occurs in May or June at White Mountain I, and in June or July at White Mountain II. Average duration of snow cover is about 160 days at the lower station and 210 days at the higher locations, but there is great variation from year to year, with a range at White Mountain II of 292 days in 1973 to 54 days in 1964. Snow depths at upper elevations generally increase until March or April, and occasionally May. Maximum recorded depth at White Mountain I is 94 in (239 cm), in March 1969, and 123 in (312 cm) at White Mountain II, in the same month. A measurement of 144 in (366 cm) was made with a snow sampler at Chiatovich Flats, at 10,600 ft (3,230 m), in March 1967. It is difficult to measure accurately such parameters as snowfall, snow cover, and snow depth; significant differences can occur in short horizontal distances. Much depends on the location and exposure of the site, the instruments used, the times of observation, and the knowledge, persistence, and hardiness of the observer. It is commonly difficult to differentiate snow that has fallen directly from the sky and that subsequently removed or deposited by wind. Thus, the figures used here should be taken as approximations.
At the base of the mountains snowfall is much lower than above 10,000 ft (3,050 m), but individual storms can still bring impressive 24-hour totals. All of the lowland stations have had daily accumulations up to 20 in (60 cm). Snow cover in the lowlands is discontinuous, with durations rarely longer than six weeks. There is notably rapid disappearance of snow cover on the western slopes of the range, facing Owens Valley, up to elevations of 8,000–9,000 ft (2,440–2,745 m), even following major snowfalls. This slope faces a high angle of the sun during the relatively warm temperatures of afternoon. As expected, snow cover lasts longer on north- and east-facing slopes
than on south- or west-facing ones — often many weeks longer. In late spring and early summer, the snowline is distinctly lower and the cover more continuous in the mountains as viewed from Fish Lake Valley in the east than from Owens Valley in the west.
Precise measurements are nonexistent, but it is obvious that some snow is removed from its site of original fall on the extensive summit upland to slopes and canyons adjacent to the plateau. Because prevailing winds generally have a westerly component, the wind-blown snow is generally deposited on eastern slopes, although occasionally the direction is reversed, especially in conjunction with northeast winds following the passage of storms or with closed cyclones in late winter and spring. Quite probably, the sites of deepest snow accumulation in the White Mountains are at the heads of the major canyons on the east side of the range. Maximum Pleistocene glaciation occurred in these canyons, the heads of which are now steep-walled, east- or northeast-facing cirques — favorable locations for the accumulation and retention of snow. A conspicuous sight from Fish Lake Valley is a discontinuous, but commonly prominent, line of snow cornices at the edge of the summit plateau, commonly lasting through the summer months and sometimes well into fall or until the next season's snowpack begins. A worthwhile addition to the knowledge of the White-Inyo Range climate would be the acquisition of reliable data on how much snow is moved by wind and deposited either east or west of the summit upland. It may well be less than visual inspection, perhaps affected by the discomfort and poor visibility caused by fine blowing snow, indicates. Sporadic measurements of wind-blown snow in the Sierra Nevada in similar terrain do not substantiate the notion of a significant increase in snowpack in leeward sites, except in localized areas.
Wind is even more difficult to measure accurately than snow. Instruments are commonly of doubtful accuracy, and there are wide fluctuations in speed and even direction in short periods of time at any one location, and across very short horizontal distances. Topography obviously exerts a major influence on wind speed and direction. Wind data, somewhat incomplete, exist for the two mountain stations, and there is a more complete and accurate record at Bishop.
Direction is easier to measure than speed. At both mountain stations, for about two-thirds of the year the prevailing direction (from which the air is moving) is westerly — northwest, southwest, and west. East is the least common, although the northeast can be of importance in some years. At Bishop prevailing directions are northerly and southerly, reflecting the topographical influence of Owens Valley. Fish Lake Valley and Deep Springs Valley very likely have the same regime as Owens Valley.
Maximum wind velocities have approached or exceeded 100 mph (45 m/s) at both mountain stations. The strongest winds usually come in the winter, from west to south in association with storm fronts, or from west to north after frontal passage. Monthly average maximum speeds are about 30 mph (13.5 m/s) during the winter
at White Mountain II and near 20 mph (9 m/s) in the summer; averages are about 5 mph (2.5 m/s) less in all seasons at White Mountain I. At Bishop average speeds are much less than at high elevation, but a peak gust of 75 mph (33.5 m/s) was recorded in August 1976. In general, maximum velocities in the lowlands occur in winter and early spring. Experience indicates that Fish Lake Valley and the eastern slopes of the White Mountains have the highest velocities of the lowland areas.
The intensity, frequency, and duration of high-velocity winds in the White Mountains do not seem extraordinary for a mountain range of its height and latitude. However, an extensive portion of the range above treeline has no shelter from wind. The most notable aspect of wind at higher elevations is not high velocity, although that may occur, but rather the constancy of moderate wind, with a conspicuous lack of calm, even in summer. Significant results of wind are the redistribution of snow, poor visibility from blowing snow, and the wind chill factor. At 10°F (-12°C), a wind speed of 30 mph (13.5 m/s) is calculated to have roughly the same effect on humans as an equivalent temperature of -33°F (-36°C) without wind.
Wind also significantly affects snow texture. At upper elevations the snow surface is seldom smooth or powdery but is generally hard-packed, ridged, and of unequal depth, with patches of bare ground present even in wet years, because of frequent wind action. The snow is commonly packed and ridged as it strikes the ground; most snowstorms occur with moderate to high wind velocities. Despite increasing attention in recent years to the White Mountains as an area for cross-country skiing and winter snow camping, snow conditions in the range are not particularly favorable for such activities. Added to the adverse surface texture is the problem that, in many years, the amount of snow accumulation is insufficient to cover ground irregularities adequately.
The traveler or resident in the White Mountains should be aware of troublesome or possibly hazardous weather and climatic conditions that can occur. Snowstorms, sometimes heavy and often accompanied by moderate to high winds, can happen from fall through spring, and even in summer. These may bring poor visibility and drifting snow, which can make vehicular, foot, or air travel difficult or dangerous. From October through May, temperatures well below freezing make frostbite a problem if adequate clothing is not worn. The wind chill factor commonly has the effect on humans of lowering the effective temperature 20 to 40°F (10 to 22°C).
June through September is the thunderstorm season, which can bring flash floods in canyons, such as the Narrows below Westgard Pass, and common lightning strokes. Anyone who has experienced a major thunderstorm above treeline in the White Mountains with little or no protection available is acutely aware of danger from lightning.
At any elevation, but particularly above 10,000 ft (3,048 m), solar radiation — especially ultraviolet — may be intense, resulting all too quickly in painful sunburn and chapped lips. Water is not readily available throughout most of the range, par-
ticularly along the road to the two mountain stations, and dehydration due to the wind and low humidity can be a problem.
Travelers into the White Mountains — by foot, ski, road vehicle, or helicopter — are strongly advised to check weather reports before entering the range. San Francisco, Los Angeles, Reno, and Salt Lake City are major forecast centers, and the National Weather Service station at Bishop can be contacted during daylight hours for briefing on local conditions. Mountain weather is notorious for rapidly developing adverse events, but knowledge of seasonal patterns, empirical observation of cloud sequences, and current information from forecast offices can do much to reduce potential hazards.
Goodridge, James D. 1981. California rainfall summary, monthly total precipitation, 1849–1980 . California Dept. of Water Resources.
Holmboe, J., and H. Klieforth. 1957. Investigations of mountain lee waves and the air flow over the Sierra Nevada, final report . Contract AF 19 (604)-708. Meteorology Department, University of California, Los Angeles.
Houghton, J. 1969. Characteristics of rainfall in the Great Basin . Desert Research Institute, University of Nevada System, Reno.
Houghton, J., C. Sakamoto, and R. Gifford. 1975. Nevada's weather and climate . Nevada Bureau of Mines and Geology, University of Nevada, Reno.
National Oceanic and Atmospheric Administration. Climatological data (1950–1990): California, vols. 54–94; Nevada, vols. 65–105 (monthly and daily tabulations and annual summary). NOAA, National Climate Center.
Neiburger, Morris, James G. Edinger, and William D. Bonner. 1982. Understanding our atmospheric environment , 2d ed. W. H. Freeman, San Francisco, p. 453.
Powell, Douglas. 1963. Physical geography of the White Mountains, California-Nevada. M.A. thesis, Geography Department, University of California, Berkeley.
White Mountain Research Station. Unpublished observations from White Mountain Research Station weather recording sites I and II. White Mountain Station, Bishop, Calif.
Deborah L. Elliott-Fisk
Geomorphology is the scientific study of landforms. This discipline, which has traditionally been a part of both geography and geology, focuses on (1) describing the various landforms that make up the earth's natural landscape, (2) determining the processes that have shaped these landforms, and (3) reconstructing the environments in which these features formed. Because of its high diversity of bedrock types, large elevational gradient, and geologic history, the White-Inyo Range possesses a diverse suite of landforms.
What is a landform? It is simply a part of the landscape that has a distinctive shape or morphology, a unit that can be delineated either qualitatively by visual means in the field or quantitatively through an analysis of morphology, composition, and relative position on the landscape. Hills, valleys, mountain peaks, dunes, and floodplains are all landforms. Landforms in the White-Inyo Range are listed by their origin in Table 2.1.
It is easiest to think of a landform as being a function of process, materials, and time (Gregory, 1978). The three external processes shaping landforms are (1) weathering, (2) erosion, and (3) deposition. These three processes will each be discussed briefly.
Weathering is the chemical decomposition and mechanical disintegration of rock materials. We can think of this as the wearing away of earth materials through time. Different weathering processes have relative degrees of importance in different environments and are partially functions of climate. Many mechanical weathering processes require temperature and moisture fluctuations. These changes exert a mechanical stress on the rock (such as frost wedging), which causes it to disintegrate along joint planes or between individual mineral grains (Fig. 2.1). Chemical weathering, like all chemical processes, is a function of temperature. With the presence of water, at higher temperatures chemical weathering will take place more rapidly.
Thus, we see different types and rates of weathering of particular rock types (such as granite) in low-, middle-, and high-elevation climates of the White-Inyo Range. Weathering rates in the range have been quantified by Denis Marchand (1968, 1970, 1971, and 1974) in his studies of weathering and soil development at Sage Hen Flat and the Cottonwood Basin. He estimated that weathering and accompanying erosion (removal of weathered materials) of 0.3 to 1.2 in (1 to 3 cm) of material occur in 1,000
years. In conjunction with this, Valmore LaMarche (1968) has measured weathering and erosion on slopes inhabited by Bristlecone Pine (Pinus longaeva ). Tree roots are exposed through time as material is removed around them. Using dendrochronology (tree-ring dating), LaMarche was able to quantify the amount of material removed for a given unit of time. Elliott-Fisk (1987) has investigated the weathering of glacial till boulders deposited by a series of glaciers over time; her results show that weathering is slow but progressive in the White Mountains. All of these researchers have shown that weathering is a relatively slow process in the dry climate of the White Mountains. Because precipitation (available moisture) increases with altitude (to the crest), weathering may be thought to be more rapid at higher elevations, but it must be remembered that temperature decreases with altitude as well, and frost and ice are weathering agents.
It may prove fruitful to look at the relative degree of soil development as an indication of weathering along an altitudinal gradient through the range. Soils form as the result of weathering and the decomposition of organic matter and are functions of climate, organisms, parent material, relief, and time (Jenny, 1941). Thus, one must try to hold these factors constant to assess the relative importance of climate (or altitude) to soil formation and weathering. It is possible to find the same geologic formation (for example, a particular granitic pluton) at different altitudes in the range. The relief, or slope position, can also be held constant. If we can determine
that the surface of the landscape has not been exposed to disturbance through various erosional or depositional agents, time can also be held constant. However, it has been difficult to locate a sequence of landforms of the same age that span the range's entire elevational gradient. Elliott-Fisk (1987) has studied soils formed on glacial deposits of different ages (i.e., allowing time to vary) and shown that soils do progressively develop through time. It is also difficult to hold the organisms, especially the vegetation, constant as one goes up an altitudinal gradient. Sagebrush (Artemisia spp.) communities span virtually all altitudes of the range, but, as their productivity varies, they do not allow soil development to be evaluated only as a function of climate.
Erosion, the second geomorphic process, is the removal of weathered material from a slope. In order for material to be removed from a slope, its initial inertia must be overcome by the mass and momentum of the erosional agent. Erosional agents include running water (streams, rivers, sheetflow, soil water, and groundwater), glacier ice, ground ice, waves, tides and currents, wind, gravity, and organisms (especially humans). Geomorphic agents are linked to climate because some of them occur only in particular climatic settings, and their relative importance varies with the climate of an area. Waves, tides, and currents are obviously limited to water bodies, and gravity is common to all slopes, regardless of climatic setting. However, running water, glacier ice, ground ice, wind, and organisms (to some degree) vary with climate. All of these terrestrial geomorphic agents have operated in the White-Inyo Range during recent geologic time.
It is possible to reconstruct climatic events of the recent or distant past by studying landforms. Beaty (1963, 1968, 1970, and 1974) has worked for many years on determining the role that debris flows play in the development of alluvial fans flanking the White Mountains (Fig. 2.2). Debris flows are triggered by intense summer thunderstorms or very rapid snowmelt, where a large amount of weathered material is catastrophically eroded, transported down valley, and deposited at the valley floor on the alluvial fan. Debris flows and other deposits that result from flooding can be devastating on the landscape below, especially if it is inhabited, so these events are regarded as natural hazards. By mapping and dating debris flows, one may be able to calculate the periodicity of the events (with changes in frequency and magnitude of the flows an indication of climatic change) and use this information to estimate the probability of their recurrence.
Another erosional agent that has left its mark on the eastern slope of the White Mountains is the glacier (Fig. 2.3). A glacier is a moving body of ice; ice actually flows through a glacier much as water slowly flows down a stream channel. The moving ice can erode and transport materials, later depositing them down gradient (e.g., down valley). A series of valley glaciers formed south to north from Wyman Creek to Trail Creek along the eastern slope of the White Mountains, on high plateaus along the crest, and in the upper reaches of a few of the western drainages (Elliott-
Fisk, 1987). Studies of glacial landforms (Elliott-Fisk, 1987; Elliott-Fisk and Dorn, 1987; Swanson et al., 1988) show at least six stages of Quaternary glaciation that correlate with the glaciations of the Sierra Nevada and also provide more detailed information on the glacial history of the region. One can reconstruct climate during these glacial events using several techniques, perhaps most directly by mapping the glacial deposits, reconstructing the glacier that deposited them, and then determining the climatic conditions necessary to form such a glacier.
After the materials are transported by the erosional agent, they are eventually deposited at a new location. Deposition is the third geomorphic process. These depositional materials commonly form distinct landforms that are attributed to a particular geomorphic agent. These landforms can then be identified as fluvial (running water), glacial (glacier), colluvial (gravity), periglacial (ground ice), aeolian (wind), marine (waves, tides, and currents), or biological (organisms) in origin. They too can then yield climatic information, as these depositional agents vary with climate.
Depositional landforms commonly can be identified by their distinctive shapes, but landforms deposited many years ago may have been worn down by weathering processes and lost their once-distinct appearance. It is still possible to identify these landforms as to their depositional agent based on knowledge of how various geomorphic agents shape and sort materials. Table 2.2 lists the characteristic shapes and degrees of sorting of sediments by various agents. This table serves as a guide to identifying depositional landforms in the White-Inyo Range. It can be seen from this chart that glacial clasts (particles) are commonly subangular and poorly sorted, in contrast to fluvial clasts, which are rounded and well sorted. (These differences are due to the densities and velocities of the fluids [agents], among other factors.) Thus, if we find a deposit with no characteristic morphology on a ridgetop, valley floor, or intermediate slope, we can commonly identify its agent of deposition. Early Quaternary high-altitude ridgerop (interfluve) deposits on Cottonwood Plateau (Fig. 2.4) and Chiatovich Flats (Fig. 2.5) have been identified as glacial in origin by sediment analyses, as surface expression of glacial landforms is lacking.
Geomorphologists have long been interested in the presence of apparent high-altitude planation surfaces at crestal and mid-elevation positions in many mountain ranges (see Fig. 2.5). These surfaces are commonly present on fault-block mountains and have been hypothesized to have been eroded in valley floor positions, then uplifted with the range. They are typically mantled with depositional material. It is possible to identify multiple planation surfaces at several distinct elevations in a particular mountain range. If these surfaces were formed at valley bottom locations, determination of their time of formation could shed light on rates of tectonic activity for a range (e.g., when was the range uplifted?) (Curry, 1984). However, many workers believe that these high-altitude surfaces were simply eroded in place by various agents, especially frost action (ground ice); as such, they have been referred to as cryoplanation surfaces. A third group of scientists believe that these surfaces were very likely the result of both past lower-elevation and present higher-elevation processes.
High-altitude erosion surfaces are distinct in the White Mountains. The best example is Pellisier Flats, which extends from an elevation of 12,400 ft (3,775 m), near Mt. Hogue at the south, to 13,400 ft (4,080 m) at The Jumpoff, just south
of Boundary and Montgomery peaks to the north. What appears to be an extremely flat surface from distant positions to the west or east of the range is actually a gently rolling surface with small, residual bedrock outcrops (inselbergs, tors, and monadnocks) (Fig. 2.6). The surface is mantled with frost-shattered debris and is presently a periglacial landscape. Both active and relict patterned ground occur here (Mitchell, LaMarche, and Lloyd, 1966; Elliott-Fisk, 1987; Mahacek-King et al., 1987) and dominate the surficial geology of other high plateaus in the range.
Ongoing research focuses on how extensively these plateaus were covered with glacial ice, which requires either the formation of an ice cap, under present topographic constraints, or the presence of valley glaciers prior to tectonic uplift and an episode of major stream incision (valley formation). Many of the high plateau surfaces are mantled by glacial deposits that have been somewhat reworked by frost action and fluvial processes (Elliott-Fisk, 1987). Chiatovich Flats (10,200–11,600 ft, or 3,130–3,560 m, elevation) and the Cottonwood Plateau (11,200–12,000 ft, or 3,435–3,680 m, elevation) are the best examples. Other surfaces are mantled by rock glaciers (such as the North Fork Perry Aiken Creek, Fig. 2.7) and felsenmeer (i.e., frost-shattered boulder fields; see Figs. 2.1 and 2.6).
Geomorphologists have long debated whether felsenmeer indicates the absence of glaciation. This is a difficult question because our knowledge of high-altitude weather-
ing rates is poor. The thin ice cover of a cold-based glacier (frozen to the surface) may allow the preservation of felsenmeer. The plateau between Mt. Barcroft and White Mountain Peak apparently had such an ice cover in the past, as large, frost-shattered blocks on the plateau above the cirques of the North and South Forks of McAfee Creek show surface polish and abrasion, indicating that this area was in the zone of accumulation of a former glacier (Elliott-Fisk, 1987). The same or an adjacent glacier may have deposited the granodiorite erratics on the dolomite surface of the Cottonwood Plateau to the south (see Fig. 2.4). If the time of formation of these glaciers can be deduced, the rate of high-altitude weathering since its disintegration can be calculated.
The tectonic activity of the earth can be characterized as an internal geomorphic process. Displaced, faulted, and warped landforms attest to local tectonic activity. Good examples of these exist along the western and eastern escarpments of the White-Inyo Range in the form of displaced alluvial fan (Fig. 2.8) and lake bed materials.
Alluvial fans are triangular or cone-shaped masses of debris at the base of a mountain front (see Fig. 2.2). Deposition of alluvial and colluvial material results when the lower gradient and wider channel of the valley floor are encountered as the material is transported down the mountain flanks. The velocity of the fluid drops, causing
the deposition of these materials. Streams spread out as distributaries at this point, instead of coming together in the form of tributaries to a main stream channel, as they do in the drainage basin above.
As one ascends the stream canyons at the apex (head) of these alluvial fans, older fan materials (fanglomerates) may be seen perched on the walls many meters above the present fan surface (as along Indian Creek, Milner Creek, and Black Canyon). It is also possible to find an older fan apex at a higher elevation back in the mountain front (as at Jeffrey Mine Canyon on the west slope of the White Mountains). If these materials can be dated, the amount of tectonic uplift can be calculated. Unfortunately, these deposits are difficult to date, as they are usually beyond the age range of radiocarbon or lack organic materials necessary for radiocarbon dating. However, along Black Canyon in the southwestern White Mountains (see Fig. 2.8), a volcanic deposit is interbedded with fanglomerate. This material is tephra (air-filled lapilli and ash) from the catastrophic eruption of the Long Valley caldera and is referred to as the Bishop Tuff (volcanic ash) (Bateman, 1965). It has been dated by uranium fission and other techniques, with the best estimate of its age currently at 0.79 Ma (millions of years before present). This deposit is present in Black Canyon 490 ft (150 m) above the valley floor, attesting to 7.6 in (19 cm) of uplift per 1,000 years.
Near the Westgard Pass road at the Waucoba Embayment of the western front of the White-Inyo Range, a series of lake beds (the Waucoba Lake beds) is exposed.
Walcott (1897) speculated that these lake deposits have been uplifted 3,300 ft (1,000 m) above the valley floor. Although intensive studies of these deposits have not yet been conducted, their age is estimated at 3 million years, suggesting an uplift rate of 13 in (33 cm) per 1,000 years.
The relatively slow tectonic uplift of the range may not seem important in reference to our short lifetimes, but some tectonic activity, generating earthquakes and displacement along fault lines, can be abrupt. Escarpments along fault lines cutting through alluvial fans suggest rapid vertical displacements of several meters. Further research needs to focus on dating this tectonic activity, which can be accomplished through studies of soil formation on the fans. The presence of ash from the Mono and Inyo craters interbedded with uplifted fan sequences and other deposits (Fig. 2.9) suggests tectonic disturbance in the last 5,000 years (Mahacek-King et al., 1987; Mahacek-King et al., 1988). It is now well known that the entire region is tectonically, and hence geomorphically, very active, as has been shown by intensive studies in the Long Valley caldera region and along the White Mountains fault zone, and by historical records of earthquake occurrence. It is believed by some workers (Curry, 1984) that this tectonic activity is accompanied by volcanism and basin formation and is migrating north from the Owens Valley (which witnessed a catastrophic eruption about 790,000 ka (thousands of years before the present) to the Mono Basin (which is still volcanically active) and the Bridgeport basin. This may be related to major plate rifting along the axis of the Gulf of California and to crustal extension.
A wide variety of materials is available for landform generation in the White-Inyo Range. All three basic rock classes (igneous, metamorphic, and sedimentary) are present and range from Precambrian to Quaternary in age (see Chapter 3). The mineralogy and petrology of these rocks vary, with some more susceptible to jointing, chemical weathering, and other processes. Thus, some types of landforms tend to be associated with distinct rock formations. For example, the Campito Formation (especially the Andrews Mountain Member), a sandstone or quartzite, is very susceptible to frost action, resulting in the formation of angular felsenmeer seen as stone stripes on slopes along White Mountain Road. Bateman (1965) contrasts this with the Montenegro Member shale, which forms slabs and breaks down rapidly to clay-rich soils. Differences in patterned ground derived from metamorphic and granitic materials are apparent at high elevations.
The White-Inyo Range is largely lacking in organic-rich deposits and landforms, in the form of either peat or fossilized organic remains. This is most likely due to the present aridity of the range and the lack of topographic depressions along the steep mountain flanks. Organic-rich deposits may compose distinct landforms, but they possess other advantages in that they are a superb source of material for the study of fossil plant and animal communities and for dating of geomorphic events (see Fig. 2.9). Recently, turf and earth hummocks have been found in the Alpine Zone of Pellisier Flats and register geomorphic change with the deposition of volcanic tephra and changing hydrology of the site (Mahacek-King et al., 1987). The arid climates of the range have allowed woodrat middens to be preserved in a diversity
of environments. Analysis of plant macrofossils from these middens is providing valuable information on the late Quaternary geologic and climatic history of the range (Jennings, 1988; Jennings et al., 1989).
Reference has been made throughout this section to the role of time in the control of landform development. If we think of the White-Inyo Range as a dynamic mountain system that is continuing to evolve, we can ask not only what landforms are developing now, but which ones were formed in the past, when they were formed, and what this tells us about the evolution of the range.
The geologic evolution of the range is discussed in Chapter 3 of this book. However, in a general sense, it can be stated that all landforms visible today are the
products of late Cenozoic processes. We can imagine many of the high peaks (such as White Mountain Peak) forming over millions of years through weathering and erosion. It is very likely that the peak was always exposed to the subaerial environment, perhaps a nunatak above a hypothetical White Mountain ice cap. Other landforms are more recent, forming with the glaciation of the range or by recent fluvial activity.
Bateman, P. C. 1965. Geology and tungsten mineralization of the Bishop District California . U.S. Geological Survey Professional Paper 470.
Beaty, C. B. 1963. Origin of alluvial fans, White Mountains, California and Nevada. Annals, Association of American Georgarphers 53:516–535.
Beaty, C. B. 1968. Sequential study of desert flooding in the White mountains of California and Nevada . Technical Report 68-31-ES. U.S. Army natick Laboratories, Natick, Mass., January.
Beaty, C. B. 1970. Age and estimated rate of accumulation of an alluvial fan, White Mountains, California, U.S.A. American Journal of Science 268:50–77.
Beaty, C. B. 1974. Debris flows, alluvial fans, and a revitalized catastrophism. Zeitshrift für Geomorphologie , N.F. Suppl. Bd. 21:39–51.
Curry, R. R. 1984. Mountain summit glacial tills and their tectonic implications, eastern Sierra Nevada, California. Abstracts, Annual Meeting of the Geological Society of America , Reno, Nevada, p. 481.
Elliott-Fisk, D. L. 1987. Glacial geomorphology of the White Mountains, California and Nevada: Establishment of a glacial chronology. Physical Geography 8:299–323.
Elliott-Fisk, D. L., and R. I. Dorn. 1987. Pleistocene glaciation of the White Mountains, CA-NV, and correlation with the Sierra Nevada. Geological Society of America, 1987 Annual Meeting, Abstracts with Programs , p. 655.
Gregory, K. J. 1978. A physical geography equation. National Geographer 12:137–141.
Jennings, S. 1988. Late Quaternary vegetation change in the White Mountain region. In C. A. Hall, Jr. and V. Doyle-Jones (eds.). Plant biology of eastern California . Natural History of the White-Invo Range, symposium vol. 2, pp. 139–147. University of California, Los Angeles.
Jennings, S. A., D. L. Elliott-Fisk, T. W. Swanson, and R. I. Dorn. 1989. A late-Pleistocene chronology of the White Mountains, CA-NV. Association of American Geographers Program Abstracts, Baltimore . Washington, D. C.
Jenny, H. 1941. Factors of soil formation . McGraw-Hill, New York.
LaMarche, V. C., Jr. 1968. Rates of slope degradation as determined from botanical evidence, White Mountains, California . U.S. Geological Survey Professional Paper 325-I, pp. 341–377.
Mahacek-King, V. L., J. A. Onken, D. L. Elliott-Fisk, and R. L. Bettinger. 1987. Quaternary silicic tephras in the White Mountains, CA-NV: Depositional environment and geomorphic history. Geological Society of America, Annual Meeting, Abstracts with Programs, p. 756.
Mahacek-King, V. L., D. L. Elliott-Fisk, T. E. Gill, and T. A. Cahill. 1988. Elemental analysis by PIXE applied to tephrochronology of the White Mountains, California-Nevada. Geological Society of America, Abstracts with Programs , vol. 20, no. 7, p. A54.
Marchand, D. E. 1968. Chemical weathering, soil formation, geobotanical correlations in a portion of the White Mountains, Mono and Inyo Countries, California. Ph. D. thesis, University of California, Berkeley.
Marchand, D. E. 1970. Soil contamination in the White Mountains, eastern California. Geological Society of America, Bulletin 81:2497–2505.
Marchand, D. E. 1971. Rates and modes of denudation, White Mountains, eastern California. American Journal of Science 270:109–135.
Marchard, D. E. 1974. Chemical weathering, soil development, and geochemical fractionation in a part of the White Mountains, Mono and Inyo Counties, California . U.S. Geological survey Professional Paper 352-J.
Mitchell, R. S., V. C. LaMarche, and R. M. Lloyd. 1966. Alpine vegetation and active frost features of Pellisier Flats, White Mountains, California. American Midland Naturalist 75:516–525.
Swanson, T. W., D. L. Elliott-Fisk, R. I. Dorn, and F. M. Phillips. 1988. Quaternary glaciation of the Chiatovich Creek Basin, White Mountains, CA-NV: A multiple dating approach. Geological Society of American, Abstract with Programs , vol. 20, no. 7, p. A209.
Walcott, C. D. 1897. The post-Pleistocene elevation of the Inyo Range, and the lake beds of Waucobi Embayment, Inyo County, California. Journal of Geology 5:340–348.
Geologic History of the White-Inyo Range
Clemens A. Nelson, Clarence A. Hall, Jr., and W. G. Ernst
Introduction and General History
The White-Inyo Range (Fig. 3.1), representing the westernmost range of the Basin and Range structural province, extends for 110 mi (175 km) from Montgomery Pass south-southeastward to Malpais Mesa opposite Owens Lake. Its maximum width, east of Bishop, is approximately 22 mi (35 km). The terminological separation of White from Inyo mountains is placed along the Westgard-Cedar Flat-Deep Springs Valley Road, a division that has no particular topographic or geologic significance. As is typical of the ranges of the province, it is bounded, generally on both sides, by normal faults of large-magnitude slip. The northern part of the range is mainly an easterly tilted block marked by an impressive escarpment from the Owens-Chalfant-Hamill valleys on the west, at 4,300 ft (1,310 m), to the crest of the range at White Mountain Peak, 14,246 ft (4,342 m). The southern, Inyo Mountains, part of the range has been tilted slightly to the west, with its maximum relief from Saline Valley at 1,100 ft (353 m) on the east to the range crest at Mt. Inyo, 11,107 ft (3,385 m).
Rocks in the White-Inyo Range span the time from the late Precambrian (700 Ma [Ma, millions of years ago]) to the Holocene, or Recent (i.e., last 10,000 years). Figure 3.2 is a simplified geologic timetable showing major time units and millions of years before the present (Ma) for the beginning of each, and the time spans of mountain building episodes (orogenies) referred to in the text. All periods of the Paleozoic (570–225 Ma), Mesozoic (225–65 Ma) and Cenozoic (65 Ma to the present) are represented, some incompletely. Rocks from the late Precambrian to the end of the Devonian Period (345 Ma) are entirely of sedimentary origin, having been deposited as sand, shale, dolomite, and limestone along the western edge of the North American continent by stream systems flowing westward into a marine basin called the Cordilleran geosyncline (Fig. 3.3). The total accumulation of uppermost Precambrian through Lower Jurassic surficial deposits in the geosyncline exceeded 4.5 mi (7 km) in thickness in the White-Inyo region. Beginning in Mississippian time, about 345 million years ago, the sedimentation pattern changed in response to elevated lands lying to the north and possibly west, resulting in the accumulation of coarse-grained sands and conglomerate, reflecting higher-energy stream systems. This pulse of uplift has been considered by some geologists as the result of the first of several collisions of the North American lithospheric plate and an ancient Pacific plate; the inferred plate collision is held responsible for the Antler orogeny, or time
of mountain building. The Pennsylvanian and Permian (320–225 Ma) were times of renewed deposition of limestone, reflecting a return to conditions of quiet carbonate bank formation.
This lull was again succeeded by extensive tectonic activity, apparently resulting from a second plate collision, the Sonoma orogeny (230–220 Ma). This event, as in
the Sierra Nevada to the west, resulted in extensive volcanism in a marine environment, the products of which are exposed as interlayered lava flows, ash beds, and continental sedimentary rocks in the southern Inyo Mountains and in the northern White Mountains.
In both the White-Inyo Range and the Sierra Nevada, surficial volcanism and its deep-seated counterpart, igneous intrusion, continued intermittently until approximately 155 million years ago as a consequence of the most intense orogeny to affect the region, the Nevadan (see Fig. 3.2). This complex episode, the result of terrane
accretion (plate collision) and amalgamation of far-traveled subcontinental fragments outboard of the North American plate and large-scale subduction of an ancient Pacific plate beneath the North American plate, produced major compressional deformation, metamorphism, and the emplacement of numerous coarse-grained intrusive bodies, termed plutons. The latest Mesozoic (90–75 Ma) was a time of renewed plutonic activity, represented by cross-cutting granitic intrusive rocks that transect the earlier-formed deformational structures. The physical conditions inferred to have attended crystallization of the metamorphic minerals associated with the intrusive rocks document the presence of a volcanic-plutonic arc marking the western margin of the North American continent throughout Mesozoic time (Fig. 3.4).
The early and middle Cenozoic Era was a time of large-scale uplift and extensive erosion; no sedimentary record of this time has been left in the White-Inyo Range. A major erosion surface truncates all previously formed rocks, including the deep-seated granitic plutons, which had solidified at depths of perhaps 6 mi (10 km) or more. Beginning approximately 10 million years ago, the range experienced the outpourings of extensive fragmental volcanic ejecta (volcanic tuff) and basaltic lava flows. Remnants of this episode can be seen in the volcanic rocks that mantle the northeast corner of the White Mountains at Montgomery Pass and east, in the table-like flows extending from the Cottonwood Basin (SE of Mt. Barcroft) to the north end of Deep Springs Valley, in the large expanse of volcanic rocks covering the saddle area between Eureka and Saline valleys on the east flank of the Inyo Mountains, and in the Malpais Mesa at the south end of the Inyo Mountains.
The latest Cenozoic was a time of renewed uplift along Basin and Range normal faults flanking the range, and transcurrent motion along the Furnace Creek fault zone at the margin of the White Mountains and Fish Lake Valley. Both styles of faulting reflect a major episode of crustal extension that was initiated in the Basin and Range Province approximately 15 million years ago. That this style of deformation continues to the present in the White-Inyo region is attested to by such events as the Chalfant earthquakes of July 1986, which measured up to 5.5 on the Richter scale.
An additional result of the uplift in latest Cenozoic time was the onset of Quaternary glaciation, which affected the area east and north of Mt. Barcroft and White Mountain Peak. This was followed by further uplift and the development of extensive aprons of alluvial deposits at the western and eastern margins of the range.
The Central White-Inyo Range
The area of principal interest for this guide extends from directly north of White Mountain Peak south to approximately 37°00¢ N. lat., a distance of 47 mi (75 km). All of this area has been mapped geologically at a scale of 1:62,500 or 1:24,000 (Fig. 3.5), and a more detailed discussion of its sedimentary, deformational, and igneous history is possible. This is also the area covered by the enclosed geologic map and by the principal and subsidiary road logs included at the end of this chapter.
The central White-Inyo Range exposes the best stratigraphic sections of the uppermost Precambrian to middle Paleozoic strata in the range. The basal part of the section (Fig. 3.6) contains the Precambrian-Cambrian transition. The Lower Cambrian portion of this section, the Waucoban Series, is regarded as the North American-type succession for rocks of this age. It contains the oldest trilobite faunas in the Americas, abundant archeocyathans (primitive reef-forming animals), numerous criss-crossing tracks and trails of primitive molluscs and arthropods, and molluscan body fossils (Wyattia and others), now regarded as marking the beginning of the Paleozoic Era. This section of strata and its fossils illustrate a remarkable explosion of life at the Precambrian-Cambrian boundary, in which "explosion" all invertebrate phyla are represented. These fossil representatives occur in the limited span extending from the upper Wyman Formation (animal tracks and trails), through the Reed and Deep Spring formations (primitive molluscs and fossil animal trails), to the Campito and Poleta formations (trilobites). This succession represents a group of rocks deposited in largely tidal and subtidal environments, as well as reefal and off-reef carbonate bank and shoal environments. The terrigenous strata are largely shale-siltstone and quartzite deposited on the shallow continental shelf of the time. It has been speculated that one could have walked across the early Paleozoic sea in the White-Inyo region in water only chest-deep. The strata contain abundant shallow water indicators, such as sedimentary rocks with cross-bedding, current and wave ripple marks, mud cracks, and very highly bioturbated beds.
The geologic structure of the area from 37°00¢ N. lat. north to just south of Mt. Barcroft is relatively simple: the range is largely anticlinal. North of Westgard Pass, the White Mountains are principally a gently south-plunging, asymmetrical anticlinorium (east flank nearly vertical) exposing the oldest stratigraphic unit, the Wyman Formation, in the core. The central, relatively simple structure (Fig. 3.7), is modified on both the west and east sides by more complex, closely appressed sets of compressional structures. The Inyo Mountains are dominated by a more open and nearly horizontal anticline with a southeast trend, also exposing the Wyman Formation at its core and similarly modified on its flanks by folds as well as faults.
Lying between the Inyo anticline and the White Mountain anticlinorium is a structural downwarp containing many closely spaced anticlinal and synclinal folds and associated faults, along the western side of Deep Springs Valley. These are the famed Poleta folds, an area used for instructional purposes by many academic departments of geology.
The Inyo Mountains, and areas to the southeast, expose structures interpreted as the consequence of extreme compression: a system of low-angle reverse faults — extensive, gently inclined surfaces across which old rocks have been thrust from beneath, onto and over younger rocks. In the Inyo Mountains, this fault system has been termed the Last Chance thrust, along which Precambrian and lower Paleozoic rocks have been juxtaposed above rocks of middle and late Paleozoic age. In its northernmost exposures, the fault is inclined to the north, suggesting that the northern Inyo Range and a part of the White Mountains may lie entirely above the Last Chance thrust.
A roughly similar interpretation can be made for the area of the White Mountains north of Mt. Barcroft. Even though this part of the range has been invaded by many Mesozoic intrusive bodies (to be discussed later), the gross distribution of the pre-intrusive strata suggests that the lower to middle Mesozoic core of the range has been overthrust by Paleozoic and Precambrian strata, which appear to lie structurally above, along the northwest, north, and northeast flanks of the range.
Dating the many structural features of the White-Inyo Range is difficult because of the large temporal gap between the deformed and undeformed strata. Many of the structures are truncated by one or more of the Mesozoic intrusive bodies. Consequently, the major folding and thrust faulting can be regarded as pre-intrusive (pre-Nevadan), generally before 180 Ma, possibly consequences of the Antler (» 350–340 Ma) and/or the Sonoma (» 230–220 Ma) mountain building episodes (see Fig. 3.2).
The uppermost Precambrian to middle Mesozoic strata of the central White-Inyo Range are essentially "swimming" in a sea of middle to late Mesozoic intrusive rocks. The plutons have disrupted pre-intrusive structures, have metamorphosed the sedimentary strata to slate, schist, and quartzite, and in some cases have strongly deformed and stretched sedimentary layers. The plutons have been radiometrically dated and range in age from 180 to 75 Ma.
The plutons of the range exhibit a variety of emplacement mechanisms. This is well illustrated by the several bodies in the southern White Mountains and the northern Inyo Mountains (see Fig. 3.7). The oldest, the Marble Canyon composite pluton and the Eureka Valley and Joshua Flat plutons, are Jurassic in age. They have invaded the axial portion of major synclinal (downwarp) and basinal structures to the east and northeast of the White Mountain and Inyo anticlines, respectively. They are in part discordant to the original sedimentary layering but have principally shouldered aside their host rocks and locally thinned them to approximately 40% of the initial thicknesses.
The slightly younger Jurassic Beer Creek and Cottonwood plutons occupy the same synclinal downwarp but are more broadly discordant, as they cut across the eastern limb of the White Mountain anticlinorium. The youngest Jurassic pluton, the Sage
Hen Flat (145 Ma), is unique among the White-Inyo intrusive rocks in that it has not significantly deformed its host rocks. It has passively invaded the west flank of the White Mountain anticlinorium, apparently by a process of magmatic stoping: blocks of the overlying host rocks were fractured and, because of their greater density, sank into the magma chamber, allowing the molten granitic column to rise buoyantly.
Cretaceous intrusive activity in this area is represented by the Birch Creek and Papoose Flat plutons, each of which attained its present structural position by a process of forcible injection. Both appear to have initially invaded their host rocks along pre-intrusive faults (see Fig. 3.7). The Birch Creek Pluton occupies the east flank of the central portion of the White Mountain anticlinorium. During emplacement, it deflected to the northwest and overturned the central anticlinal structure and the strata of the east flank. The Papoose Flat Pluton was emplaced within the southwest limb of the Inyo anticline, drastically disrupting the southeast trend of the structure, producing a pronounced westward bulge and overturning the strata along its northeast border. Accompanying the bulging, the sedimentary succession at the contact with the pluton was in places stretched and thinned to less than 10% of initial thicknesses, an extreme example of forcible emplacement and accompanying attenuation of wall rocks.
North of the area of Fig. 3.7, Jurassic plutonism is represented by the Barcroft Pluton (see enclosed geologic map). The pluton transects the White Mountain anti-clinorium, diverting the western limb of the structure from its regional northward trend to the northeast, parallel to the southeast border of the intrusive (Fig. 3.8). Alternatively, this change in trend could be related to reverse movement along the Marmot thrust, a south-dipping fault within the Precambrian-Cambrian rocks directly south of the Barcroft Pluton. It has also been suggested that the pluton was emplaced along a steeply inclined structural discontinuity along which the Paleozoic rocks of the southern block had been uplifted or overthrust into contact with the Mesozoic rocks of the northern block.
The White Mountain anticlinorium with its N-S trend is thought by some to have developed during the Sonoma orogeny of latest Paleozoic to early Mesozoic age. The Cottonwood Pluton (Fig. 3.8) intruded preexisting folds in the Precambrian and Cambrian sequence and is apparently responsible for the overturning of the east limb of the anticlinorium.
The dominant minor fold axes in the Wyman and Deep Spring formations are approximately N-S and are nearly horizontal. Exposure of the Wyman and Deep Spring formations in the northeast part of Fig. 3.8 and south of the Marmot thrust are more highly deformed than in the southern part of the area depicted in Fig. 3.8. The dominant trends of minor fold axes in this more deformed area are N30°E. These folds are interpreted to be a result of deformation associated with the Marmot thrust and postdate the folds with N-S trends. An alternative interpretation is that the more highly deformed Wyman and Deep Spring formations in this area are related to emplacement of the Barcroft Pluton, and that the Marmot thrust follows the preexisting structural grain or was reactivitated during intrusion of the granodiorite pluton.
Folds with NE-SW or E-W trends in the Poleta Formation (see Fig. 3.8) north of the Marmot thrust are associated with the emplacement of the Barcroft Pluton. Metamorphic minerals (tremolite, scapolite, and diopside) in the Wyman Formation of this area and south of the surface trace of the Marmot thrust could reflect either the close proximity of the Barcroft Pluton at depth or, less likely, metamorphism that occurred during movement along the Marmot thrust. Thinning of the Reed Dolomite and the Deep Spring Formation in the northeast part of the area and above the Wyman-cored anticline or dome, as depicted in Fig. 3.8, suggests either: regional upwarping during deposition of the Reed and Deep Spring formations; thinning related to deformation associated with the White Mountain anticlinorium, along whose trend the dome in the Wyman Formation lies; or attenuation owing to the proximity of the Barcroft and Cottonwood plutons in a manner similar to that proposed for the Papoose Flat Pluton. Hence, folding or upwarping within the study area could have first occurred during the Cambrian or late Precambrian, and again during the Antler and/or Sonoma orogenies, or perhaps partly during the Nevadan tectonic events.
Plutons of Cretaceous age (90–75 Ma) in the region to the east and northeast of Mt. Barcroft and White Mountain Peak are represented by the Indian Garden Creek and McAfee Creek granitic bodies, which occupy large areas on the east slope of the White Mountains (see enclosed geologic map). The Indian Garden Creek Pluton is entirely within and discordant to the Cottonwood Pluton. The McAfee Creek Pluton intrudes Precambrian strata, the Cottonwood Pluton, and the Barcroft Pluton.
Recrystallization (metamorphism) of the Precambrian, Paleozoic, and Mesozoic strata of the White-Inyo Range took place repeatedly in response to the periods of heating and deformation just described. The early metamorphic mineral assemblages were overprinted and were partly to completely replaced by other assemblages produced by the younger thermal events. The rocks are typified by a regional development of greenschist facies minerals; common phases in recrystallized, stratified rocks include quartz, albite, microcline, white mica, chlorite, biotite, epidote, and magnetite. Thermal annealing reflects more intense baking occasioned by the emplacement of plutons of various ages; many, but not all, have produced distinct metamorphic aureoles or halos in the surrounding strata, indicated by the localized formation of new, higher-temperature minerals such as garnet, cordierite, tremolite, tourmaline, fluorite, scapolite, andalusite, sillimanite, and calcic plagioclase.
In the central White-Inyo Range, the record of the early to middle Cenozoic (65-15 Ma) is missing. During this time, broad regional uplift resulted in the unroofing of the range, exposing the deep-seated granitic bodies. Erosional products of the uplift and resulting denudation were probably shed chiefly westward, across what is now the Sierra Nevada (then low-lying), to the Cenozoic seaway in the region of the present San Joaquin Valley. With the development of crustal extension beginning about 15 million years ago, erosional debris was transported to adjacent down-dropped basins such as the Owens, Fish Lake, Eureka, and Saline valleys.
The character of the erosional surface produced by this long period of denudation is well illustrated in the area of Cottonwood Basin, southeast of Mt. Barcroft. From there to east of the north end of Deep Springs Valley, Miocene volcanic tuff and
overlying basaltic flows (10 Ma) lie upon a generally smooth surface planed across the Jurassic plutonic rocks of the range. This erosional surface and its basaltic cover was uplifted and tilted, mainly to the east, as shown by its elevation of 11,800 ft (3,595 m), 2.5 mi east of Mt. Barcroft on the northwest, to 5,500 ft (1,675 m) in Deep Springs Valley, from where it has been uplifted along a marginal fault to an elevation of 7,700 ft (2,350 m) east of the valley at Piper Mountain.
The development of Deep Springs Valley is typical of the late Cenozoic history of the region. The valley lies wholly within the range and trends northeastward across the regional structural grain. It is marked by a major fault system along its eastern margin, as illustrated by the steep escarpment, with a relief of 2,400 ft (730 m), east of the valley, and numerous small fault scarps, which attest to the recency of uplift. Geophysical data suggest that structural relief between the granitic rocks on the east side of the valley to the granite surface below the valley fill is as much as 5,000 ft (1,520 m) and that the alluvial fill beneath Deep Springs Lake on the east side of the valley is approximately 2,600 ft (790 m) thick.
An additional aspect of the crustal extension that the region has experienced is illustrated by the fault system that marks the margin between the northeast flank of the White Mountains and Fish Lake Valley. This is the Furnace Creek fault, which extends from Death Valley to its termination at the northwest end of Fish Lake Valley. It is a major transcurrent fault, along which the White Mountain block has been moved, principally horizontally, northwestward perhaps as much as 13 mi (20 km) relative to rocks on the northeast side of the fault.
As discussed previously, uplift during the late Cenozoic resulted in elevations sufficiently high to support small valley glaciers during Quaternary time. Small cirque basins and glacial moraines occupy the upper reaches of Leidy, Perry Aiken, McAfee, and Cottonwood creeks on the east slope of the White Mountains.
A major, relatively recent geological event, the extrusion of the Bishop Tuff from the Long Valley caldera east of Mammoth Lakes, had a marginal effect on the White Mountains. The event, which took place 700,000 years ago, was a huge volcanic explosion that resulted in the outpouring of pumice and ash amounting to more than 144 cu mi (600 cu km) and the collapse of the base of the caldera to below sea level. The volcanic tuff resulting from this eruption mantles or underlies much of northern Owens Valley and most of Chalfant and Hamill valleys, and it occurs interbedded with alluvial materials flanking the west slope of the White Mountains and in several canyons, notably Black Canyon, southeast of Bishop.
That uplift and associated volcanism has continued to the present is suggested by the presence of the very recent scarps along the range margins, especially marked along the east side of Deep Springs Valley (Road Log C), by a series of young volcanic cinder cones on the west side of the Inyo Mountains southeast of Big Pine, and by the historic earthquakes of the general eastern Sierra Nevada-Owens Valley region. This region has had a long and diverse geologic history, and geologic activity continues to the present day.
Road Logs A, B, C
Clemens A. Nelson
with botanical additions
by Mary DeDecker and James Morefield
In the following road logs, figures in the left column are distances between points of interest with comments; figures in the next column (in parentheses) represent cumulative mileage.
Frequent reference is made to clock directions to various features, such as prominent peaks (e.g., Mt. Tom is at 11:00). The clock is oriented with noon straight ahead.
For each road log, reference is made to appropriate U.S. Geological Survey 15-min-topographic quadrangle maps showing road networks.
Road Log A
Owens Valley Laboratory (OVL) to Barcroft Station of White Mountain Research Station (WMRS) and return (Bishop, Big Pine, Waucoba Mountain, Blanco Mountain, and Mt. Barcroft 15-min quandrangles).
Road Log B
Owens Valley to Papoose Flat (Big Pine and Waucoba Mountain 15-min quadrangles).
For travel to Papoose Flat, either a 4 ́ 4 vehicle or one with extra-low gear is essential. Log begins at intersection of Waucoba Road and Westgard Pass Road.
Road Log C
Cedar Flat to Deep Springs Valley (Blanco Mountain and Soldier Pass 15-min quandrangles).
Bateman, P. C. 1965. Geology and tungsten mineralization of the Bishop District, California. U.S. Geological Society Professional Paper 470.
Ernst, W. G., and C. A. Hall. 1987. Geology of the Mount Barcroft-Blanco Mountain area, eastern California. Geological Society of America Map and Chart Series. Map MCH066.
Hanson, R. B. 1986. Geology of Mesozoic metavolcanic and metasedimentary rocks, northern White Mountains, California. Ph.D. dissertation, University of California, Los Angeles.
Oakeshott, G. B. 1978. California's changing landscapes: A guide to the geology of the state . McGraw-Hill, New York.
Rinehart, C. D., and W. C. Smith. 1982. Earthquakes and young volcanoes, along the eastern Sierra Nevada . Genny Smith, Palo Alto, Calif.
Ross, D. C. 1965. Geology of the Independence quadrangle, Inyo County, California. U.S. Geological Survey Bull. 1181-O.
Sharp, R. P. 1976. Geology field guide to southern California , rev. ed. Kendall Hunt, Dubuque, Iowa.
Smith, Genny, ed. 1978. Deepest valley: Guide to Owens Valley . Wm. Kaufman, Los Altos, Calif.
U.S. Geological Survey Geologic Maps
Crowder, D. F., P. T. Robinson, and D. L. Harris. 1972. Geologic map of the Benton quadrangle, Mono County, California, and Esmeralda and Mineral Counties, Nevada. U.S. Geological Survey Map GQ-1013.
Crowder, D. F., and M. F. Sheridan. 1972. Geologic map of the White Mountain peak quandrangle, Mono County, California. U.S. Geological Survey Map GQ-1012.
Krauskopf, K. B. 1971. Geologic map of the Mt. Barcroft Quandrangle, California-Nevada. U.S. Geological Survey Map GQ-960.
McKee, E. H., and C. A. Nelson. 1967. Geologic map of Soldier Pass quandrangle, California and Nevada. U.S. Geological Survey Map GQ-654.
Nelson, C. A. 1966. Geologic map of the Waucoba Mtn. quadrangle, Inyo County, California. U.S. Geological Survey Map GQ-528.
Nelson, C. A. 1966. Geologic map of the Blanco Mtn. quadrangle, Inyo and Mono Counties, California. U.S. Geological Survey Map GQ-529.
Nelson, C. A. 1971. Geologic map of the Waucoba Spring Quadrangle, Inyo County, California. U.S. Geological Survey Map GQ-921.
Robinson, P. T., and D. F. Crowder. 1973. Geologic map of the Davis Mountain quadrangle, Esmeralda and Mineral Counties, Nevada, and Mono County, California. U.S. Geological Survey Map GQ-1078.
Ross, D. C. 1967. Geologic map of Waucoba Wash quadrangle. U.S. Geological Survey Map GQ-612.
Stewart, J. H., P. T. Robinson, J. P. Albers, and D. F. Crowder. 1974. Geologic map of the Piper Peak quadrangle, Nevada-California. U.S. Geological Survey Map GQ-1186.