INTRODUCTION
California and Nevada compose Segment 1 of the Ground Water Atlas of the United States. Segment 1 is a region of pronounced physiographic and climatic contrasts. From the Cascade Mountains and the Sierra Nevada of northern California, where precipitation is abundant, to the Great Basin in Nevada and the deserts of southern California, which have the most arid environments in the United States, few regions exhibit such a diversity of topography or environment. 
Since the discovery of gold in the mid-1800¹s, California has experienced a population, industrial, and agricultural boom unrivaled by that of any other State. Water needs in California are very large, and the State leads the United States in agricultural and municipal water use. The demand for water exceeds the natural water supply in many agricultural and nearly all urban areas. As a result, water is impounded by reservoirs in areas of surplus and transported to areas of scarcity by an extensive network of aqueducts. 
Unlike California, which has a relative abundance of water, development in Nevada has been limited by a scarcity of recoverable freshwater. The Truckee, the Carson, the Walker, the Humboldt, and the Colorado Rivers are the only perennial streams of significance in the State. The individual basin-fill aquifers, which together compose the largest known ground-water reserves, receive little annual recharge and are easily depleted. Nevada is sparsely populated, except for the Las Vegas, the Reno­Sparks, and the Carson City areas, which rely heavily on imported water for public supplies. Although important to the economy of Nevada, agriculture has not been developed to the same degree as in California due, in large part, to a scarcity of water. Some additional ground-water development might be possible in Nevada through prudent management of the basin-fill aquifers and increased utilization of ground water in the little-developed carbonate-rock aquifers that underlie the eastern one-half of the State. The potential problem of withdrawals in excess of natural recharge, however, will require careful management of ground-water withdrawals.
Climate
 The diverse physiography and north-south extent of Segment 1 result in marked climatic contrasts within the region. Five climate types in the Segment are based primarily on differences in temperature and rainfall (fig. 1): 
€ Temperate Oceanic‹Adequate precipitation in all seasons, moderate summers and mild winters, cloudy conditions prevail
€ Highland‹Altitude, for the most part, controls the weather, large amounts of precipitation as rain and snow in the mountains (fig. 2), large diurnal temperature ranges, rain shadows present on the leeward sides of mountain ranges
€ Mediterranean (Subtropical Dry Summer)‹Modest precipitation in winter, warm summers and mild winters, abundant sunshine
€ Steppe (Semiarid)‹Little precipitation, falling mostly in winter, large annual temperature range
€ Desert (Arid)‹High temperature, scant precipitation, large diurnal and annual temperature ranges, low relative humidity, little cloud cover

Precipitation and Temperature 
In California, much of the climatic variation results from the patterns of global weather systems. Precipitation is greater in the northern part of the State than elsewhere (fig. 3). However, prominent mountain ranges in California and western Nevada also have an important influence on moisture distribution in the region. Temperatures are cooler in the higher altitudes of the mountains. As eastward-moving, moist, unstable air masses rise up the western slopes of the mountains, the air is cooled and water vapor condenses and falls as rain, snow, or ice. When these air masses descend the eastern slopes, they become warmer and more stable and thus retain most of the remaining moisture. Consequently, precipitation amounts are much greater on the western slopes of the north-south-trending mountain ranges of western and eastern California, whereas semiarid to arid conditions prevail east of the mountains, as in much of Nevada and in central and southern California. 
Storms that bring moisture to the region are most frequent in winter; about 80 percent of the annual precipitation falls between October and April. The extreme northern part of California has slightly wetter summers than the rest of the segment. However, amounts of precipitation vary greatly from year to year; for example, from 1860 to 1980 the average annual precipitation of Sacramento, Calif., was 18 inches, but precipitation ranged from 35 to 195 percent of the average annual precipitation (fig. 4), or from about 6 to 35 inches per year. Fog occurs frequently on the coast and provides some additional moisture that is used primarily by vegetation.
Mountain ranges that parallel the coast also affect temperature distribution. Seaward of the mountains, temperature is moderated by the ocean, and the range between daily high and low temperatures is usually less than 20 degrees Fahrenheit. Winters are cool, but they are not generally cold in coastal areas, although temperatures drop sufficiently in the coastal part of northern California to cause some frost and a dormant season for plants. Summers in coastal areas are mild, but temperatures occasionally become hot in southern California. In contrast, the valleys east of the coastal mountains experience much greater temperature extremes. In these valleys, summer daytime temperatures can be greater than 90 degrees but fall to 55 degrees or less at night. Winters in the interior valleys are relatively mild, and freezes are uncommon. 
 Temperature ranges in the mountains of western California and eastern Nevada, as well as in the desert parts of Nevada and southern California, are much greater than in other parts of Segment 1. In the mountains and deserts very little moisture is in the air to absorb the rays of the sun, or to retain heat at night. Consequently, solar radiation is intense during the day, but the heat stored in the ground is released rapidly after sunset. Temperature extremes are hotter and colder in the desert than in other lowland areas elsewhere in the segment, but mountainous areas are warm in summer and extremely cold in winter.
 
Runoff
Runoff is the amount of water left from precipitation that can be measured as streamflow after losses to evaporation, transpiration by plants, and the replenishment of storage within the aquifers. The areal distribution of runoff from 1951 to 1980 (fig. 5) closely followed the areal distribution of precipitation for the same period (fig. 3), but the relative amounts of runoff varied as a result of climatic conditions. Runoff is greatest in the mountains, where the majority of precipitation falls as snow, which melts in the spring and runs off with minimal evapotranspiration (the process by which liquid water is converted to water vapor either by evaporation or by transpiration from plants). Runoff is greater than 40 inches per year in many mountainous areas. The basins in the arid parts of Nevada and southeastern California have virtually zero runoff because most precipitation that falls is evaporated almost immediately. However, high-intensity storms or rapid snowmelt in the mountains that border the basins may cause flash floods that reach the floors of the basins. Coastal areas have a direct relation between the amount of precipitation and runoff. 

Water Surplus and Deficit
 The relation between precipitation and evapotranspiration is a major factor in water availability. Generally, if annual precipitation exceeds annual potential evapotranspiration, then there is a net surplus of water and streamflow is perennial. However, annual potential evapotranspiration can exceed annual precipitation, which causes a net deficit of water. A net annual moisture deficit is present almost everywhere in California and Nevada (fig. 6). The only areas with an annual moisture surplus are the northern California coast, which receives considerable rainfall from winter storms, and the mountainous regions of northern and east-central California, where condensation of water vapor in rising, moist air masses results in abundant rain or snow. Water is available to recharge aquifers only at times when precipitation or snowmelt is greater than actual evapotranspiration. Thus, not all areas mapped as having a net water surplus in figure 6 are recharge areas. 
In most of Nevada and in southern California, nearly all streams that head in the mountains are ephemeral and lose flow to alluvial aquifers within a short distance of where the streams leave the mountains and emerge onto the valley floors. In much of northern California and in the Humboldt, the Truckee, the Carson, and the Walker River drainages of Nevada, however, runoff is sufficient to support perennial streams. The Colorado River is supplied primarily by runoff from the Rocky Mountains. Before the inception of agriculture, the largest rivers in the vast Central Valley of California overflowed their banks during periods of peak winter flows and formed extensive marshlands. An elaborate flood control system and the lowering of the water table by withdrawals for irrigation now keep these rivers within their banks.
The geographical distribution of moisture in Segment 1 greatly influences patterns of agricultural and urban development. Much of Nevada receives little precipitation, and, consequently, ground- and surface-water supplies are limited. This limitation has severely restricted urban development and has put constraints on agricultural development. Las Vegas, the largest urban area in the State, obtains most of its water from the Colorado River, which is many miles away.
 California receives relatively abundant precipitation. However, the precipitation is concentrated in areas of the State remote from most of the large urban centers and major agricultural areas. A further complication is the unpredictability of precipitation on an annual basis, which can often make surface-water supplies undependable. To provide a dependable, year-round supply of water to areas where it is most needed is a full-time, massive undertaking, and is accomplished by careful water management and an extensive water-transportation network (fig. 7).
Physiography and land use
	The physiography of the region (fig. 8) is a product of the geologic history of the area.  Several coastal mountain ranges underlain by severely folded, faulted, commonly metamorphosed marine and continental sediments, form the Pacific Border and the Lower Californian Physiographic Provinces.  In the interior, the granitic rocks that underlie the fault blocks of the Sierra Nevada and the volcanic rocks of the southern Cascade Mountains join to form the eastern border of the low-lying California Trough, which contains the Central Valley.  East of the Sierra Nevada, the land is characterized by a series of low, north-south-trending mountain ranges and intervening valleys (fig. 9); the ranges and valleys were created by faulting that resulted in the horst and graben structures that form the Basin and Range Physiographic Province.  In the extreme northeastern part of Nevada, the southernmost extent of the Columbia Plateaus Physiographic Province is formed by basalt lava flows.
	Land use in Segment 1 is directly related to topography and the availability of water. Major land uses in California and Nevada are shown in figure  10. The flat floor of the Central Valley of California, one of the Nation¹s most important agricultural areas, is used almost entirely for growing crops. Most of the cropland, however, must be irrigated. The mountains that surround the Central Valley are areas of rugged topography and, accordingly, are used predominantly as forest and woodland, even though they receive large amounts of precipitation. Almost all of Nevada and  large parts of southern California receive little precipitation; accordingly, most of the land in these areas is desert shrubland (compare figs. 5 and 10), although sufficient water is available to allow livestock  to be grazed in some places.
	The major cities in the coastal areas of California appear as large areas of urban sprawl on figure 10. Although coastal California receives moderate to large amounts of precipitation, surface-water and ground-water supplies in those urban areas are not sufficient to provide the water needs of the population. As a result, a huge network of reservoirs, canals, and aqueducts has been constructed in California to transport water to these urban areas and other areas of water deficit.


MAJOR AQUIFERS 
Ground water in Segment 1 is contained in five major aquifers (fig. 11), four of which consist primarily of basin-fill deposits that occupy structural depressions caused by deformation of the Earth¹s crust. The four basin-fill aquifers are the Basin and Range aquifers, the Central Valley aquifer system, the Coastal Basins aquifers, and the northern California basin-fill aquifers. The fifth major aquifer is the northern California volcanic-rock aquifers. Few of these aquifers extend over an area large enough to be termed ³regional.² An exception is the Central Valley of California, which is a very large basin-fill aquifer best described as a ³regional² aquifer. Some water in Segment 1 is obtained from areally-extensive volcanic and carbonate rocks, but water within these rocks is mostly in fractures or solution openings and, consequently, the rocks generally yield little water. One notable exception is in eastern Nevada, where some alluvial basins are hydraulically connected by widespread deposits of permeable carbonate rocks that underlie the alluvium. In some places, consolidated rocks are hydraulically connected to overlying unconsolidated deposits and, thus, are part of the same aquifer or aquifer system. 
Because of the scattered, local nature of most of the aquifers, it is difficult to classify them. The grouping used herein is based on geology, physiography, and climate but is not the only one possible. The areas designated ³not a principal aquifer² lack sufficient basin-fill sediments or permeable consolidated rock to yield significant amounts of water to wells.
 The Basin and Range aquifers are located in an area that comprises most of Nevada and the southern California desert (fig. 11). The water-yielding materials in this area are in valleys and basins, and consist primarily of unconsolidated alluvial-fan deposits, although locally flood plain and lacustrine (lake) beach deposits may yield water to wells. Also, the consolidated volcanic and carbonate rocks that underlie the unconsolidated alluvium are a source of water if the consolidated rocks are sufficiently fractured or have solution openings. Many of these valleys and basins are internally drained; that is, water from precipitation that falls within the basin recharges the aquifer and ultimately discharges to the land surface and evaporates within the basin. Ground water is generally under unconfined, or water-table, conditions at the margins of the basins, but as the unconsolidated deposits become finer grained toward the centers of the basins, the water becomes confined. Rarely, basins might be hydraulically connected in the subsurface by fractures or solution openings in the underlying bedrock. These multiple-basin systems end in a terminal discharge area, or sink, from which water leaves the flow system by evaporation. Also, several basins or valleys may develop surface-water drainage that hydraulically connects the basins, and ground water flows between the basins, mostly through the unconsolidated alluvial stream/flood plain sediments. 
 The Central Valley aquifer system (fig.11) occupies most of a large basin in central California between the Sierra Nevada and the Coast Range Mountains. The Central Valley is the single most important source of agricultural products in the United States, and ground water for irrigation has been essential in the development of that industry. The basin contains a single, large, basin-fill aquifer system, the largest such system in the Nation. Although the valley is filled with tens of thousands of feet of unconsolidated sediments, most of the fresh ground water is at depths of less than 2,500 feet. Ground water in the valley is under unconfined to confined (artesian) conditions, primarily depending on depth; most of the shallow ground water is unconfined.
 The Coastal Basins aquifers occupy a number of basins in coastal areas from northern to southern California (fig. 11). These basins have similar morphology and a Mediterranean climate. All are in structural depressions formed by folding and faulting, all are filled with marine and alluvial sediments, and all are drained by streams that contain water at least part of the year. Nearly all the large population centers in Segment 1 are located in these basins, and the available ground water is used primarily for municipal supplies. In most of the basins, however, population has grown to such an extent that local ground-water supplies are no longer adequate, and surface water must be transported from distant sources to meet demand. Ground water in the basins is under unconfined to confined conditions, and two or more vertically sequential aquifers can be present in a basin, separated by confining units that consist of fine-grained sediments. In nearly all basins that contain more than one aquifer, however, the aquifers are hydraulically connected to some degree. Seawater intrusion is a common problem in nearly all the Coastal Basins aquifers.
 Interior northern California is sparsely populated, and most ground-water demand is for agricultural irrigation. The most productive and highly-utilized aquifers in the area are the northern California basin-fill aquifers (fig. 11). These aquifers are in unconsolidated alluvial sediments. However, in some basins, wells drilled into underlying volcanic rocks might produce large quantities of water, often more than wells completed in the unconsolidated sediments.
 The northern California volcanic-rock aquifers consist of volcanic rocks that yield water primarily from fractures and locally from intergranular spaces in porous tuffs. Because water-yielding zones in these rocks are unevenly distributed, wells that yield water are outnumbered by dry holes; however, in some areas, wells completed in the volcanic-rock aquifers yield large volumes of water. The northern California volcanic-rock aquifers are relatively unexplored and undeveloped. 

GEOLOGY AND GEOLOGIC HISTORY
Rocks and deposits exposed at the surface in Segment 1 range in age from Precambrian to Quaternary (fig. 12). They consist of igneous intrusive rocks, pyroclastic and extrusive volcanic rocks, and marine and continental sediments, many of which, particularly the older rocks (pre-Mesozoic) have been intensely metamorphosed, folded, and faulted. 
The principal water-yielding units are unconsolidated continental clastic deposits of Cenozoic age that partly fill structural basins created by faulting. Volcanic rocks, which are principally lava and pyroclastic flows of Cenozoic age, are important aquifers in scattered areas. Paleozoic limestones and dolomites associated with basin-fill Cenozoic clastic deposits in eastern Nevada are the only older rocks with significant water-yielding potential.
During Precambrian time and the Paleozoic Era, an almost uniform thickness of approximately 40,000 feet of marine sediments was deposited in the Cordilleran geosyncline. This geosyncline was an elongated trough that extended north to south in western North America and included the area that is now eastern Nevada and southern California. Sedimentation was marked by two periods of alternating clastic and carbonate deposition that resulted in the following sequence: quartzite and siltstone, limestone and dolomite, argillite and quartzite, and limestone.
At the end of the Paleozoic Era, volcanoes were active on a grand scale in eastern California and western Nevada. This volcanism marked the beginning of igneous activity that was to become increasingly important during Mesozoic time.
A number of shallow marine invasions inundated parts of the region during the Mesozoic Era. Conditions were such that marine formations alternated with nonmarine deposits derived from erosion of rocks in the continental interior. During this time, a coastal strip as much as 400 miles wide was formed by a combination of marine sedimentation and igneous activity, granitic intrusions, and subaerial volcanism, and was welded to the western margin of the preexisting continental mass.
The early Mesozoic seas spread inland as far as central Utah and Wyoming but were soon blocked by a narrow uplift in central Nevada. During the remainder of the Mesozoic Era, only intermittent subaerial deposition took place east of this uplift. West of the uplift, a thick sequence of Mesozoic marine and continental sediments was deposited, interspersed with lava flows, volcanic breccia, and tuff.
The close of middle Mesozoic time culminated in the first great orogeny in the western part of North America since Precambrian time. As mountain ranges rose, the marine, continental, and volcanic deposits of the Pacific Coast were folded, metamorphosed, and complexly faulted. Intense deformation of the older rocks spread eastward across most of Nevada. Late in the Mesozoic Era, the Pacific coastal region was again downwarped and the sea intruded.
During the Cenozoic Era, volcanic rocks and sedimentary deposits accumulated over wide areas of Segment 1, to thicknesses of as much as 50,000 feet. Early in the Cenozoic Era, the Basin and Range area was a high mountain surface with external drainage. During middle to late Cenozoic time, however, large-scale block faulting formed the Coast Range Mountains, the California Trough, and the Sierra Nevada and caused the Basin and Range structures. These structures are a sequence of alternating horsts and grabens that trend north-south and are reflected in the present-day topography. Volcanism, which still continues today, formed much of the Cascade Mountains.
In late Cenozoic time, the California Trough and the structural basins in the Coast Range were filled with marine and terrestrial deposits that ranged from a few thousand to as much as 50,000 feet in thickness. The grabens of the Basin and Range were filled with continental deposits and minor lava flows to thicknesses of generally less than 2,000 feet, but locally as much as 50,000 feet. The late Cenozoic also was the time of development of basins in the mountains of northern California and Nevada. These basins were filled with clastic sediments and numerous basaltic lava flows.

FRESH GROUND-WATER WITHDRAWALS
 Ground water is an important resource in California and Nevada and accounted for nearly 40 percent of all freshwater used in the two States during 1985 (fig. 13). Fresh ground-water withdrawals in California during this period were about 16 times as much as those in Nevada. In Segment 1, irrigated agriculture accounts for the greatest amount of ground-water use, followed by withdrawals for public supplies. More than 25 million people, or about 66 percent of the population of the two-State area, depend on publicly supplied ground water. 
Total withdrawals of fresh ground water during 1985, by county, are shown in figure 14. Counties with the largest withdrawals are those where vast areas are irrigated, such as the Central Valley of California, or counties having large population centers. The large withdrawals shown for central Nevada and some southern California counties are somewhat misleading. These areas are largely desert, and because of the extremely large size of some of the counties, small withdrawals in scattered pumping centers plot as unrealistically large withdrawals when totalled for counties that contain hundreds of thousands of square miles.
The Central Valley aquifer system had the largest ground-water withdrawal in Segment 1 during 1985 (fig. 15). Approximately 9,000 million gallons per day (about 10 million acre-feet per year) was withdrawn from the Central Valley aquifer system. Of that amount, approximately 8,000 million gallons per day, or 8.9 million acre-feet per year, was withdrawn for irrigation and accounted for about 11.5 percent of all ground-water withdrawals in the United States. One acre-foot, or 43,560 cubic feet of water, is the volume of water that will cover an area of 1 acre to a depth of 1 foot.
The Coastal Basins aquifers supply the largest population centers in Segment 1 and are second only to the Central Valley aquifer system in total ground-water withdrawals (fig. 15). Much ground water is withdrawn for agricultural use in these coastal basins, but public supply accounted for about 54 percent of the approximately 4,370 million gallons per day (about 4.9 million acre-feet per year) withdrawn during 1985; this is due primarily to the large population in the coastal cities of southern California that depend heavily on ground water for public supply.
Irrigated agriculture is the largest user of fresh ground water withdrawn from the Basin and Range aquifers. Although ground water is a significant source of water for public supplies, large population centers, such as Las Vegas, the Carson City area, and the Reno­Sparks area, depend heavily on surface water for their public supplies. The desert basins receive little precipitation during the year, and surface and ground water are scarce, which limits population growth in the region.
The northern California volcanic-rock aquifers and the northern California basin-fill aquifers together supplied only 5 percent of the total fresh ground water withdrawn in Segment 1 during 1985. These aquifers compose only a small part of the segment, and the demand for ground water in northern California is not great. 

Introduction
The part of Segment 1 east of the Southern Cascade Mountains, the Sierra Nevada, and the smaller mountain ranges east of the Los Angeles­San Diego area is called the Basin and Range Physiographic Province (fig. 16) and contains three principal aquifer types collectively referred to as the ³Basin and Range aquifers.² These aquifers underlie most of Nevada and parts of eastern and southern California, western Utah, southern Arizona, southwestern New Mexico, and southern Oregon and Idaho; their extent is approximately, but not exactly, the same as that of the physiographic province. The aquifers are formed of volcanic and carbonate rocks and unconsolidated to consolidated basin-fill deposits. The basin-fill deposits form the most productive aquifers and are generally in individual alluvial basins that are drained internally and are separated by low mountains (fig. 17). Except for small areas that drain to the Colorado River, no streams that rise within the Basin and Range Province carry water to the oceans. Practically all the precipitation that falls in the area is returned to the atmosphere by evapotranspiration, either directly from the soil or from the lakes and playas that occupy the lowest points within the basins and that are discharge areas for the alluvial aquifers.
The Basin and Range Province is the most arid area in the Nation; the potential annual water loss through evapotrans-piration exceeds the annual water gain from precipitation even at the higher elevations (fig. 18). Clear skies and low humidity cause extreme daily and seasonal temperature ranges as the sparsely covered land surface is heated quickly by solar radiation and then rapidly cools at nightfall. In more humid climates, the denser vegetative cover uses energy derived from solar radiation to drive the process of evapotranspiration, thus moderating diurnal and seasonal temperature variations.
Each of the large desert basins has an area where the land slopes toward a central depression, and each has a main drainageway that is dry most of the time. Many of the valleys have playas in their lowest depressions (fig. 19). The playas are left by the evaporation of intermittent lakes. Parts of some of the valleys have become encrusted to a depth of several inches with alkaline salts, which cover the surface as a powdery crust. However, in some valleys, permanent lakes that have no outlets are fed by surface drainage and contain saline or alkaline water, produced when dissolved minerals are concentrated by evaporation of the lake water.
Geohydrologic Units
Within the Basin and Range Province, aquifers are not continuous, or regional, because of the complex faulting in the region. Three principal aquifer types collectively called the Basin and Range aquifers in this report are volcanic-rock aquifers, which are primarily tuff, rhyolite, or basalt of Tertiary age; carbonate-rock aquifers, which are primarily limestones and dolomites of Mesozoic and Paleozoic age; and basin-fill aquifers, which are primarily unconsolidated sand and gravel of Quaternary and Tertiary age (fig. 20). Any or all three aquifer types may be in, or underlie, a particular basin and constitute three separate sources of water; however, the aquifers may be hydraulically connected to form a single source. Other rock types within the region have low permeability and act as boundaries to the flow of fresh ground water. The aquifers in the Great Basin part of the Basin and Range Province (fig. 16) were studied as part of the U.S. Geological Survey¹s Regional Aquifer-System Analysis (RASA) Program.

Volcanic-Rock Aquifers
The volcanic-rock aquifers (fig. 20) can be separated into three categories‹welded tuffs, bedded tuffs, and lava flows. The different characteristics for the storage and transmission of water in each category depend on the presence of primary and secondary porosity. Physical characteristics that affect the movement of ground water include the number and degree of interconnection of joints, the relation of joint density to degree of welding and compaction, the horizontal partings within tuffs, the development of rubble zones between lava flows, and the interconnection of vesicles in the lavas. 
Ash-flow tuffs are consolidated deposits of volcanic ash, which were emplaced by flowage of a turbulent mixture of gas and pyroclastic materials. Ash-flow deposits consist principally of glass shards and pumice fragments that are usually less than 0.15 inch in length, although some flows consist of ejecta of larger size. Typically, the deposits are nonsorted and do not exhibit bedding, in contrast with the generally pronounced bedding of ash-fall tuff deposits. In general, ash flows are tens of feet thick, but some are only a few feet thick, whereas others are hundreds of feet thick. After emplacement of an ash flow, compaction or welding of the ash can result in an average 50-percent reduction in the porosity of the original flow.
Welding within a single ash flow is variable, and each ash flow can be categorized by three distinct orders of welding‹none, partial, or dense. Commonly, a zone of dense welding is underlain and overlain by zones of partial welding, which are, in turn, underlain and overlain by zones of no welding (fig. 21). However, in some thin, exceptionally hot flows, the entire unit of tuff can be densely welded. The degree of welding directly affects the interstitial porosity of the ash-flow tuff. In the nonwelded base or top of a fresh ash flow, the interstitial porosity can be greater than 50 percent; in the densely welded part, it can be less than 5 percent. 
Columnar jointing characterizes the zones of dense and partial welding; these joints form in response to tensional forces that develop as the flow cools. Columnar-joint spacings range from a few tenths of an inch to many feet; the more closely spaced joints are usually in the zone of most intense welding. The joints are usually vertical, but departures from the vertical are common. Cooling joints are not common in the nonwelded parts of the ash flow (fig. 21). 
The joints in outcrop in the ash-flow tuffs are polygonal joints that formed as the flow cooled and other joints that formed after cooling as a result of compaction of underlying, porous, bedded tuff or from regional tectonic stresses. Both types of joints are restricted mostly to the dense, brittle, welded tuff and die out or markedly decrease within the underlying and overlying partially welded zone (fig. 21). The polygonal structure is generally obscured by the joints that formed after cooling, except in the youngest welded tuffs. Horizontal partings are locally a few tenths of an inch wide and tens of feet long. Because the partings parallel the foliation within the welded zone or the contact between flows, they may represent breakage along a plane of primary weakness after the removal of overburden; therefore, the partings are not likely to be open at depth and are limited in extent. 
The bedded-tuff aquifers are ash-fall tuffs that consist of poorly to well sorted, friable particles the size of fine sand to granules. Locally, the ash-fall tuffs either have been reworked by running water or were originally deposited in standing water. The friable nature of these rocks prevents the formation of open joints or faults within them; as an example, open fractures were not seen in hundreds of feet of tunnels dug through these rocks beneath Rainier Mesa on the Nevada Test Site near Las Vegas. Where glass shards are altered to clay minerals, the permeability of the ash-fall tuffs is reduced by several orders of magnitude. 
The lava-flow aquifers consist of basalt or rhyolite and have not been studied in detail. No laboratory determinations for porosity and permeability have been done on these aquifers because the movement of ground water through them is controlled mostly by porosity developed along cooling joints and in rubble zones between individual lava flows. Basalt flows might be a texturally heterogeneous mass that laterally and vertically ranges from congealed, dense, impermeable lava to highly porous zones that consist of loosely consolidated cinders. The texture depends, for the most part, on the amount of gas present in the lava when the flow erupted. Permeable zones, which consist of masses of basalt rubble, are at the tops of some dense lava-flow surfaces and are overlain by subsequent flows or by sediments. The dense lava flows, which have minimal primary permeability, might be fractured by regional stresses, resulting in high secondary permeability. When fracture systems interconnect with highly permeable rubble and cinder zones, the rock mass tends to be highly transmissive. 

Carbonate-Rock Aquifers
Thick sequences of Mesozoic and Paleozoic carbonate rocks underlie many of the alluvial basins in southeastern California and eastern Nevada within the Basin and Range Province; these rocks also extend into western Utah and southeastern Idaho. Results of deep drilling indicate that intervals of cavernous carbonate rock are as deep as 5,000 feet and might locally extend to depths of 15,000 feet. In some test wells, circulation of drilling fluid has been extremely difficult to maintain and, in a few, the downhole drilling equipment has suddenly dropped. Both conditions indicate that the carbonate rock is cavernous. 
Quartzite, shale, siltstone, sandstone, and some limestone and dolomite of Early Cambrian and late Precambrian age underlie the carbonate rocks in the eastern part of the Basin and Range Province. However, these rocks have minimal primary and secondary permeability, and probably form the lower boundary of the carbonate-rock aquifers. 
The carbonate-rock aquifers can be divided into two parts‹an upper rock sequence of Late Triassic to Early Mississippian age that consists primarily of limestone with minor amounts of dolomite, interbedded with shale and sandstone, and a lower sequence of limestone and dolomite of Middle Devonian to Middle Cambrian age that contains little clastic material. The total thickness of carbonate rocks may be greater than 15,000 feet, but, as a result of the combination of deep erosion and structural deformation, this thickness is rare in any one location. The saturated thickness of the carbonate strata ranges from a few hundred to more than 10,000 feet and depends on the combined influence of geologic structure, erosion, and depth to water. In general, because of the great 
aggregate thickness and stratigraphic position of the rocks that compose the carbonate-rock aquifers, several thousand feet of an individual aquifer is within the zone of saturation throughout most of the areal extent of the aquifers. Such an aquifer is completely unsaturated only in the vicinity of its outcrop area and is totally absent only atop buried structural highs. 
The carbonate rocks are highly fractured and are locally brecciated (fig. 22). Individual outcrops of the aquifers can exhibit three or more sets of joints, one or more high-angle faults, and one or more brecciated zones. For example, in the Nevada Test Site area near Las Vegas, Nev., the joints and most of the faults in the carbonate rocks are steeply inclined fractures. Brecciation commonly occurs along faults showing only a few feet of displacement and does not necessarily reflect movement of large magnitude. Joint density bears a strong relation to rock type; fine-grained carbonate rocks have the greatest joint density. Generally, the joints divide the rock into blocks that range from 1 inch to a few inches on a side. Medium-grained carbonate rocks are divided into blocks that range from a few inches to 1 foot on a side, whereas blocks of coarse-grained carbonate rocks commonly range from 6 inches to 2 feet on a side. 
In outcrop, secondary openings are locally along bedding planes in the carbonate rocks, but no widespread connection of such openings is known. Some of the bedding-plane openings might have formed entirely by subaerial mechanical and chemical weathering, but some might have formed by partial dissolution of the rock. Dissolution, presumably in the subsurface, has created small, smooth, tabular openings along otherwise tightly closed bedding and joint planes (fig. 22). 
Basin-Fill Aquifers
Before the most recent period of tectonic activity, which began in middle Miocene time (about 17 million years before present), the Basin and Range region was characterized by moderate relief, and streams in the region did not have enough power to transport large volumes of sediments. As the mountains were uplifted, however, stream gradients increased and the transporting power of the streams greatly increased. Steep, narrow canyons and gulches were incised into the sharp escarpments that bounded the mountain ranges and enormous volumes of material were eroded from the mountains. In some places, blocks of sandstone greater than 10 feet in diameter were transported several miles from their outcrop areas onto flat areas beyond the mouths of canyons. The sediments eroded, transported, and deposited by the streams are the principal material of basin-fill aquifers (fig. 23). Some of the older basin-fill deposits (Miocene and Pliocene age) are consolidated; however, the basin fill consists mostly of unconsolidated deposits of Pliocene through Holocene age.
The most permeable basin-fill deposits are present in the depressions created by late Tertiary to Quaternary block faulting and can be classified by origin as alluvial-fan, lake-bed, or fluvial deposits. At the time of major deposition, the climate was more humid than the modern climate. Lakes were in most of the closed basins and some basins were connected by streams. In general, the coarsest materials (gravel and boulders) were deposited near the mountains, and the finer materials (sand and clay) were deposited in the central parts of the basins or in the lakes. Occasionally, torrential storms produced heavy runoff that carried coarse material farther from the mountains and resulted in the interfingering of fine and coarse material. The distribution of sediment size is directly associated with distance from the mountains. Three geomorphic landforms can be distinguished on the basis of the gradient of the land surface. Alluvial fans border the mountains and have the steepest surface slopes and the coarsest sediments (fig. 24). Basinward, individual alluvial fans flatten, coalesce, and form alluvial slopes of moderate gradient. A playa, or dry lake bed with a flat surface, is present in the lowest part of the basin, usually at or near the center of the basin (fig. 19), and most of the sediment deposited on the playa is fine grained. 
The most important hydrologic features of the basins are the alluvial fans. The basin fill receives most of its recharge through the coarse sediments deposited in the fans. These highly permeable deposits allow rapid infiltration of water as streams exit the valleys that are cut into the almost impermeable rock of the surrounding mountains and flow out onto the surface of the fans. The coarse and fine sediments within the alluvial fans are complexly interbedded and interfingering    (fig. 24) because the position of the distributary streams that transported the sediments continually shifted across the top of the fan.
Material deposited in perennial lakes or in playas consists principally of clay and silt with minor amounts of sand and is present in all of the basins. In most places, these sediments include some salts deposited by evaporation. The clay and salt deposits merge laterally into coarse-grained deposits of the alluvial slopes. Minor well-sorted beach sand and gravel locally are in the subsurface near the shores of once perennial lakes. 
Fluvial deposits of Holocene age in the basins consist primarily of alluvial sand and gravel and are present along the courses of modern or ancestral streams that generally parallel the long axes of the basins. Quaternary fluvial deposits in stream channels usually exhibit a greater degree of sorting than the alluvial-fan deposits. 

Ground-Water Flow Systems
The ground-water flow systems of the Basin and Range area are in individual basins or in two or more hydraulically connected basins through which ground water flows to a terminal discharge point or sink. Except for relatively small areas that drain to the Colorado River, water is not discharged to major surface-water bodies but is lost solely through evapotranspiration. Each basin has essentially the same characteristics‹the impermeable rocks of the mountain ranges serve as boundaries to the flow system, and the majority of the ground water flows through basin-fill deposits. In the area where carbonate rocks underlie the basins, substantial quantities of water can flow between basins through the carbonate rocks and into the basin-fill deposits, but this water also is ultimately discharged by evapotranspiration. Most recharge to the basin-fill deposits originates in the mountains as snowmelt, and, where the mountain streams emerge from bedrock channels, the water infiltrates into the alluvial fans and replenishes the basin-fill aquifer. Intense thunderstorms may provide some direct recharge to the basin-fill deposits, but, in most cases, any rainfall that infiltrates the soil is either immediately evaporated or taken up as soil moisture; little water percolates downward through the unsaturated zone to reach the water table in the valleys. In mountain areas underlain by permeable carbonate rocks, most of the recharge may enter the carbonate rocks and little water remains to supply runoff. 
Because regional aquifers are not continuous within the Basin and Range area, the individual basins, which are encircled by topographic drainage divides, have been classified as one of four types based on similar recharge-discharge relations (fig. 25 and table 1). The simplest type is the ³undrained, closed basin,² a single valley in which the underlying and surrounding bedrock is practically impermeable and does not allow interbasin flow, and all recharge is discharged at a sink represented by a playa near the center of the basin. Basins underlain by permeable bedrock commonly are hydraulically connected as multiple valley systems. The ³partly drained, closed basin² is underlain or surrounded by bedrock that is moderately permeable and allows some ground water to flow out of the basin. In this type of basin, some water is evaporated or transpired at the upgradient side of a playa, but most of the water continues to flow past the downgradient side of the playa and leaves the basin. The ³drained, closed basin² has a deep water table that prevents evapotranspiration. The bedrock is sufficiently permeable to allow all recharge to flow through it and out of the basin. The ³terminal sink basin² is underlain or surrounded by bedrock that is sufficiently permeable to conduct flow into the basin, and the playa in the basin is the discharge point for recharge from several connected basins. 
In some places, an existing or ancestral stream course connects several basins that are not closed. The individual basins connected by such streams can also be classified as partly drained, drained, or terminal sink. 
Examples of the individual type basins follow, except for the drained closed basin. With the exception that the water table in the drained closed basin is far below the playa, the partly drained and drained closed basins are sufficiently similar that discussion of each type is not warranted. 

Single, Undrained, Closed Basin
Antelope Valley, Calif., which is an example of a single, undrained, closed basin, is a large topographic and ground-water basin in the western part of the Mojave Desert in southern California. Antelope Valley occupies part of a structural depression that has been downfaulted between the Garlock and the Cottonwood­Rosamond Faults and the San Andreas Fault Zone (fig. 26). Consolidated rocks that yield virtually no water underlie the basin and crop out in the highlands that surround the basin. They consist of igneous and metamorphic rocks of pre-Tertiary age that are overlain by indurated continental rocks of Tertiary age interbedded with lava flows.
Alluvium and interbedded lacustrine deposits of Quaternary age are the important aquifers within the closed basin and have accumulated to a thickness of as much as 1,600 feet (fig. 27). The alluvium is unconsolidated to moderately consolidated, poorly sorted gravel, sand, silt, and clay. Older units of the alluvium are more compact and consolidated, somewhat coarser grained, more weathered, and more poorly sorted than the younger units. The rate at which water moves through the alluvium (the hydraulic conductivity of the alluvium) decreases with increasing depth. 
During the depositional history of Antelope Valley, a large intermittent lake occupied the central part of the basin and was the site of accumulation of fine-grained material. The rates of deposition varied with the rates of precipitation. During periods of relatively heavy precipitation, massive beds of blue clay formed in a deep perennial lake. During periods of light precipitation, thin beds of clay and evaporative salt deposits formed in playas or in shallow intermittent lakes. Individual beds of the massive blue clay can be as much as 100 feet thick and are interbedded with lenses of coarser material as much as 20 feet thick. The clay yields virtually no water to wells, but the interbedded coarser material can yield considerable volumes of water. 
During deposition of the lacustrine deposits, alluvial material that was supplied from the San Gabriel Mountains encroached upon the lake and forced it northward, which resulted in a northward transgression of alluvium over lacustrine deposits. The subsurface extent of the buried lacustrine deposits is shown in figure 27. The lacustrine deposits underlie the central part of the basin and have a somewhat lenticular shape. The thickest section is near the center of the basin, and the deposits thin towards the edges of the basin. Near Little Buttes and near the east and north edges of Rogers Lake, the deposits pinch out (fig. 27, section B­B¹). Along the northern and southern boundaries of the basin, the lacustrine deposits are about 100 and 400 feet thick, respectively, where they abut buried escarpments of consolidated rocks (fig. 27, section A­A¹). Near the southern limit of the basin, southeast of Lancaster, the lacustrine deposits are buried beneath about 800 feet of alluvium, but near Rosamond Lake, they are exposed at the surface (fig. 27, section A­A¹). 
Two aquifers, which are separated by the lacustrine deposits, are in the alluvial material (fig. 27). The upper aquifer is the principal and most used aquifer and contains water under unconfined, or water table, conditions. Where the lower, or deep, aquifer underlies lacustrine deposits, it contains water under confined, or artesian, conditions; elsewhere, unconfined conditions prevail. 
Transmissivity values for the principal aquifer (fig. 28) are estimated to range from less than 1,000 to more than 10,000 feet squared per day. The transmissivity of an aquifer is a measure of how rapidly water will pass through the aquifer; the greater the transmissivity, the faster the movement of the water and the more water the aquifer will yield to wells. Where the principal aquifer is thin, either near its boundaries or on the uplifted parts of fault blocks, its transmissivity is low; where the aquifer is thick or consists of coarse-grained deposits, or both, the transmissivity is high. 
The estimated transmissivity of the deep aquifer (fig. 29) ranges from about 2,000 to 10,000 feet squared per day and is greatest where the aquifer is thick. The transmissivity of the deep aquifer varies less than that of the principal aquifer (compare figs. 28 and 29) probably because the thickness of the deep aquifer is more uniform than that of the principal aquifer.
Ground water in the Antelope Valley Basin moves from the base of the San Gabriel and the Tehachapi Mountains toward Rosamond Lake in the north-central part of the basin (fig. 30). As ground water moves eastward across the western limit of the lacustrine deposits, part of the water moves above the lacustrine deposits to recharge the principal aquifer and part moves below the lacustrine deposits to recharge the deep aquifer. Major faults that cut the alluvial deposits in Antelope Valley, especially the Randsburg­Mojave Fault (fig. 30), act as partial barriers to the movement of ground water. Water-level differences of more than 300 feet in the same aquifer are present across the Randsburg­Mojave Fault. Along several other faults, the water table is several tens of feet higher on the upgradient side of the fault than on the downgradient side. 
An estimate of the shape of the predevelopment potentiometric surface of the principal aquifer in 1915 (fig. 30) shows that before extensive pumping began, the water table was near the land surface in the central part of the basin; ground water moved northward and northeastward, and discharged by evapotranspiration at Rosamond Lake, which was dry. Withdrawal of ground water from the principal aquifer and the subsequent lowering of the water table reduced this natural discharge. By 1961, the direction of ground-water movement in the principal aquifer had been reversed from northeastward to southward and southeastward, toward the center of the basin in the area immediately southeast of Rosamond Lake (fig. 31). The main change in the potentiometric surface was the development of areas of low water levels near the withdrawal centers and the resultant reversal in the direction of ground-water flow near these areas. 
Ground water leaks through the lacustrine deposits between the principal and deep aquifers even though the lacustrine deposits do not readily yield water to wells. Based on the hydraulic heads for the two aquifers, water leaks downward along the western and southern periphery of the lacustrine deposits. In the north-central part of the area underlain by the lacustrine deposits, water leaks upward. Because of the large withdrawals from the principal aquifer, the area of upward leakage has expanded toward the areas of concentrated withdrawal (fig. 31).
The aquifers in Antelope Valley are recharged primarily by infiltration of streamflow that originates in the mountainous areas that surround the valley. The average annual precipitation on the valley floor is less than 10 inches, and runoff is minor. For the most part, streamflow that enters the valley is intermittent. During storm periods, streamflow enters the valley along its perimeter and moves across the surface of the alluvial fans toward the playas at Rosamond and Rogers Lakes. As the streams flow across the alluvial fans, all the streamflow generally infiltrates the permeable surficial deposits on the fans. Because of the desert conditions, much of the infiltrating water is quickly lost by evaporation or as transpiration by riparian vegetation. The remainder of the water infiltrates downward through the alluvial deposits until it reaches the water table. 
 The drainage area tributary to Antelope Valley is about 385 square miles. Runoff from about 20 percent of this area is measured and the collective average annual discharge at the measured points is about 24,300 acre-feet. By calculating the measured runoff per unit area and extrapolating this value to unmeasured areas, the total runoff that enters the valley was estimated to be 40,700 acre-feet. 
Evapotranspiration is the major natural discharge of ground water in Antelope Valley. Ground water generally discharges by evaporation from the water table where the water table is within 10 feet of the land surface, and, where vegetation is present, transpiration may also occur. Evaporation from an open body of water in Antelope Valley was measured at about 114 inches per year, which is an upper limit for evaporation of ground water. Because evapotranspiration from the ground-water system is complex, exact values cannot be determined. 
The use of ground water for agriculture in Antelope Valley began about 1880, when wells were drilled near the center of the valley and yielded flowing water in quantities sufficient for irrigation. In 1891, more than 100 wells were in use, but most had stopped flowing. About 1915, intense use of ground water began when a large number of wells were drilled and equipped with pumps. An estimate of annual withdrawal rates from 1915 to 1975 is shown in figure 32. The maximum rate of withdrawal of about 400,000 acre-feet per year is about 10 times the estimated annual recharge to the basin. Water removed from storage in the aquifers was a major part of the ground-water withdrawals, and severe water-level declines resulted. By about 1950, studies showed that ground-water withdrawals in the valley were greatly in excess of natural recharge and withdrawals were curtailed. 
The geographic distribution of withdrawals was generally unchanged between 1915 and 1960. After 1960, withdrawals were redistributed by abandoning some wells and adding some new wells. With the new distribution, the center of withdrawal was split into two areas; one was approximately 5 miles southeast, and the other approximately 10 miles southwest of Rosamond Lake (fig. 31). 
Withdrawals from the deep, or confined, aquifer in Antelope Valley have caused an increase in leakage to the deep aquifer from the principal aquifer along the western and southern peripheries of the lacustrine deposits. This leakage has locally lowered the water table in the principal aquifer and has resulted in the reduction of natural discharge from the aquifer. Most of the declines in the principal aquifer, however, are the consequence of withdrawals from that aquifer. Field data are not available to show the effects of water-level declines on the amount of natural discharge, but the results of a digital flow model indicate that most of the natural evapotranspiration from the center of the valley might have ceased by 1950 (fig. 33) because water levels in the principal aquifer were too deep to allow evaporation or transpiration. 
Ground water in closed basins is commonly highly mineralized because discharge by evapotranspiration increases the concentrations of minerals in the water. Some of the minerals might precipitate at or near the center of the basin. However, dissolved-solids concentrations in ground water remained practically the same or decreased slightly in Antelope Valley between 1908 and 1955 (fig. 34); this was probably caused by the reduced evapotranspiration that resulted from declining water levels in the principal aquifer.
Partly Drained, Closed Basin
The Pahrump Valley, an example of a partly drained, closed basin, covers about 1,050 square miles in Nye and Clark Counties, Nev., and Inyo and San Bernardino Counties, Calif. (fig. 35). The Spring Mountains, which form the northeastern border of the basin, are the dominant topographic feature and are the source of all the water that enters the basin. The southwestern side of the Spring Mountains is characterized by large alluvial fans that head high in the canyons that lead from Charleston Peak. The most prominent of these fans have coalesced to form the Pahrump and the Manse Fans. 
The Pahrump Valley is part of an intervalley ground-water flow system. The regional movement of ground water is generally southwestward to low areas adjacent to the Amargosa River. The major areas of ground-water discharge downgradient from the Pahrump Valley are between the towns of Tecopa and Shoshone, Calif., which are 10 to 15 miles southwest of the topographic boundary of the Pahrump Valley (fig. 35). 
Mountain-building activity in southern Nevada has affected the ground-water flow system in the Pahrump Valley. Several large thrust faults are exposed in the Spring Mountains and at the northern end of the Nopah Range (fig. 36). In some places, low-permeability clastic rocks have been displaced by the faulting so that they are above or adjacent to water-yielding carbonate rocks and thus restrict ground-water movement in the carbonate rocks. Under some conditions, permeable zones of broken rock along the fault planes might be conduits for ground water. Springs and stands of mesquite along the northwestern sides of these faults, however, suggest that the faults are barriers to ground-water flow and that the ground water moves upward along the barriers until it emerges at the land surface. Folding, associated with the faults, produced joints and fractures in some of the rocks, resulting in significant secondary permeability. 
Two distinct aquifers are in the Pahrump Valley‹the carbonate-rock aquifer, formed of carbonate rocks that bound and underlie the valley, and the basin-fill aquifer, which consists of unconsolidated deposits that have accumulated in the structural depression of the valley (fig. 37). The carbonate rocks transmit water readily and carry significant ground-water flow from the Pahrump Valley into the adjacent Chicago and Amargosa River Valleys to the southwest. The carbonate-rock aquifer is virtually undeveloped and significant future development is improbable because it is necessary to drill wells to great depths in order to obtain adequate yields. The basin-fill aquifer is, therefore, the source of virtually all withdrawals. 
The carbonate-rock aquifer consists primarily of carbonate rocks of Triassic to Cambrian age that crop out in the Spring Mountains (fig. 36) and underlie the basin fill of the Pahrump Valley (fig. 37). The aquifer extends westward and southwestward through the Nopah and the Resting Springs Ranges into the California and the Chicago Valleys (figs. 35 and 37). Because no well in the Pahrump Valley penetrates the carbonate-rock aquifer, the hydraulic properties of the aquifer are inferred from information obtained in other areas. Hydraulic continuity in the carbonate-rock aquifer is the result of an extensive network of interconnected fractures and, to a small degree, of localized solution openings. Estimates of transmissivity from nearby localities outside the valley ranged from 130 to 120,000 feet squared per day as determined from aquifer-test data from 10 wells. The greater the transmissivity, the more water the aquifer will yield. The wide range in estimated transmissivity might not be randomly distributed and might reflect variations that result from faulting as well as the number and size of solution openings.
The basin-fill aquifer consists of unconsolidated alluvial and lacustrine deposits that partly fill the structural depression of the Pahrump Valley. Coarse-grained materials have been deposited near the sides of the valley, and fine-grained lacustrine materials are in the central parts of the valley (fig. 37). The approximate areal extent of the basin-fill aquifer is 650 square miles, or about two-thirds of the total area of the Pahrump Valley. To the northeast, northwest, and southwest, the aquifer is bounded by consolidated rocks of the Spring Mountains and the Resting Springs, the Nopah, and the Kingston Ranges. To the southeast, the aquifer is bounded by a ground-water divide beneath a topographic high that separates the Pahrump and the Mesquite Valleys. Because no ground water flows across this divide, the ground-water flow systems of the two valleys are separate. 
Wells drilled into the basin-fill aquifer range from several tens of feet to more than 1,000 feet deep. With the exception of one or two wells near the margin of the aquifer, the wells do not fully penetrate the basin fill. Therefore, the thickness of the basin-fill aquifer was estimated from geophysical meas-urements. The maximum thickness of the aquifer is about 4,800 feet in the central part of the valley (fig. 38). In general, the thickest accumulations of basin fill parallel the axis of the valley. The area of maximum thickness is offset slightly toward the south end of the valley, suggesting some faulting or folding in that area. 
Estimates of the transmissivity of the basin-fill aquifer (fig. 39) are representative only of the upper 1,000 feet of the aquifer, which is the part penetrated by most wells. Variations in transmissivity are related to the deposition of the coarser materials and the position of the water table. Transmissivity values increase from the edge of the Spring Mountains, where the saturated materials are thin, toward the center of the valley, where the land surface is flatter, the water table approaches the land surface, and the aquifer is thickest. The increase in saturated thickness within the zone of coarse materials provides the highest transmissivity values; values are greater than 4,000 feet squared per day in the Pahrump and the Manse Fans. Transmissivity values decrease in nearly parallel bands across the valley to less than 1,000 feet squared per day as the sediments become finer and the saturated thickness lessens near the mountains on the southwest side of the valley. 
Virtually all the ground water in the Pahrump Valley is derived from precipitation. Most ground-water recharge occurs in the mountains, where percolating water moves through bedrock fractures to the zone of saturation, and on the upper slopes of the alluvial fans, where streamflow percolates through the unsaturated basin fill downward to the zone of saturation. The general slope of the ground-water surface in the Pahrump Valley before development (before 1913) is shown in figure 40. This map was constructed by using the earliest measurements available and shows the approximate configuration of the potentiometric surface of the basin-fill aquifer. Ground-water flow was generally from the principal recharge areas adjacent to the Spring Mountains, southwestward across the valley towards the Nopah Range. Water left the valley by evapotranspiration in the areas of shallow ground water and by subsurface outflow beneath the Nopah Range. The contours in figure 40 suggest that as ground water flowed southwest across the northwestern part of the Pahrump Valley, it moved into and through outcrops of carbonate rocks. The final discharge area for this water is not known with certainty. The 2,600-foot contour in figure 40 indicates that the hydraulic gradient in the northwestern part of the valley is toward the Ash Meadows discharge area in the Amargosa Desert north and west of the Pahrump Valley (fig. 35). However, the majority of the ground-water flow probably moved to a discharge area along the Amargosa River between the towns of Shoshone and Tecopa, which are southwest of the Pahrump Valley. 
Ground water has been developed to support agriculture in the Pahrump Valley for many years. Two large springs, Bennetts and Manse Springs (fig. 35), provided water to early travelers and were soon developed as a source of supply for irrigation. In the late 1800¹s, Bennetts Spring reportedly discharged about 7.5 cubic feet per second (5,430 acre-feet per year), and Manse Spring, about 6 cubic feet per second (4,340 acre-feet per year); most of this water was diverted to agriculture. Spring flow decreased dramatically in 1913 when ground-water withdrawals began. Bennetts Spring eventually ceased to flow as the water table declined in response to withdrawals (fig. 41), and Manse Spring ceased to flow during the 1975 irrigation season. 
Ground water provides the water supply for virtually all uses in the Pahrump Valley. The first well was drilled in the valley in 1910 in an unsuccessful attempt to obtain a flowing well. However, in 1913, three flowing wells were successfully completed, and, by 1916, a total of 28 wells had been drilled, 15 of which were flowing. The number of new wells drilled and the annual withdrawals increased slowly until the mid-1940¹s when large-capacity wells were installed. From the mid-1940¹s through 1962, the annual discharge from wells increased from an estimated 4,000 to about 28,000 acre-feet (fig. 42). 
From 1962 to 1975, population growth caused significant change in land use in the Pahrump Valley. Ground-water withdrawals increased rapidly between 1962 and 1968 (fig. 42) but decreased after 1968 as agricultural land was taken out of service and subdivided for residential use. Population of the valley increased from about 250 in 1962 to nearly 1,500 in 1975, and, when this land becomes fully developed, ground-water withdrawals could surpass those of 1968. 
Between 1913 and 1975, nearly 700,000 acre-feet of ground-water withdrawals and about 550,000 acre-feet of spring flow had been discharged from the basin-fill aquifer in the Pahrump Valley. The two most apparent effects of the discharge were large ground-water level declines and a cessation in spring discharge. Water levels have been declining since the first wells were constructed in 1913. Variations in the annual rate of decline and the net change between predevelopment and 1975 water levels (fig. 43) among different locations depend on the distribution of withdrawal, the hydraulic properties of the basin fill, and the depth of the well being measured. Hydrographs of six wells are shown in figure 44 to illustrate the typical response of water levels to withdrawals in various parts of the valley. Generally, the greatest water-level declines (about 100 feet) were along the terminal parts of the Pahrump and the Manse Fans. Two wells located in the central part of the valley, away from the major concentration of withdrawal, showed substantially less water-level decline. 
Withdrawals in the Pahrump Valley are distributed to capture ground-water flow downgradient of the Pahrump and Manse Fans and capture the discharge of Bennetts and Manse Springs in an effective manner (fig. 45). Consequently, spring discharge began to decrease shortly after withdrawals began and continued to decrease until 1975, when discharge of both springs ceased entirely during the irrigation season (fig. 41). Bennetts Spring ceased to flow in 1959; as of 1975, Manse Spring was dry during the summer irrigation season but recovered to discharge about 200 acre-feet during the winter months. The changes in spring discharge from 1875 through 1975 are shown in figure 41. The amount by which spring discharge has decreased is the amount of water captured by withdrawals. Thus, as of 1975, about 9,800 acre-feet per year of spring discharge had been captured by ground-water withdrawals. 

Terminal Sink Basin
The Salton Trough, which is a topographic and structural trough that extends from southeastern California into Mexico, is an example of a terminal sink basin. The Salton Trough is divided into two parts by the Salton Sea‹the Imperial Valley to the south and the Coachella Valley to the north (fig. 46). The Imperial Valley is the largest area of desert irrigation development in the United States. Importation of water from the Colorado River has transformed 500,000 acres of parched desert lands into one of the most productive agricultural areas in the Nation. The Coachella Valley has experienced dramatic declines in ground-water levels due to public water supply demands for a growing population, stimulated by the attraction of the Palm Springs area. 
The desert climate of the Salton Trough is characterized by extreme aridity and high summer temperatures. Average annual precipitation is slightly less than 3 inches on the valley floor and about 40 inches at the crests of the San Jacinto Mountains. Maximum summer temperatures commonly exceed 104 degrees Fahrenheit, and winter minimums are seldom below 32 degrees Fahrenheit. 
The Salton Trough is about 130 miles long and as much as 70 miles wide (fig. 46). The trough is a landward extension of the depression that is partially filled by the Gulf of California. The trough and gulf are separated by the broad fan-shaped subaerial delta of the Colorado River. Much of the land surface of the trough is below sea level and, before the delta was formed, the trough may have been part of the Gulf of California. The lowest part of the trough is occupied by the Salton Sea, whose surface is at an elevation of more than 200 feet below sea level. Most of the surface drainage is intermittent and is toward the Salton Sea in the central part of the valley. Surface water moves northwestward to the Salton Sea from the boundary between the United States and Mexico and southeastward from the San Gorgonio Pass.
The Imperial Valley occupies the wider southern part of the Salton Trough. The valley ends at the Salton Sea to the northwest and continues southward into Mexico as the Mexicali Valley. The Chocolate Mountains border the valley to the northeast, and the Peninsular Range of Baja California and southern California borders it to the southwest. The floor of the valley slopes northwestward from about sea level at the international boundary to about 230 feet below sea level at the Salton Sea. The Salton Trough was once occupied by prehistoric Lake Cahuilla. In the eastern and western parts of the trough, several ancient lake shorelines (fig. 47) are at elevations of 42 to 50 feet above sea level. 
The central part of the Imperial Valley is a large area of cultivated land entirely within the shorelines of prehistoric Lake Cahuilla. Most of the central Imperial Valley is a monotonous plain dissected by the Alamo and the New Rivers, which have incised trenches as much as 40 feet deep in soft, silty lacustrine deposits. Much of the entrenchment took place during 1905 through 1907 when virtually the entire Colorado River flowed uncontrolled in these channels and established the present-day Salton Sea. Before 1905, the center of the trough had been a playa. The lowest point of the surface of the trough is beneath the southern part of the Salton Sea and is about 275 feet below sea level. 
The Coachella Valley in the northwestern part of the Salton Trough (fig. 46) extends from the east end of San Gorgonio Pass southeastward to the Salton Sea. It is bordered on the north and east by the Little San Bernardino Mountains and on the southwest by the San Jacinto and the Santa Rosa Mountains. The Whitewater River is the main drainage in the valley.
The Coachella Valley is affected by the San Andreas Fault system, which is a complex strike-slip fault system that includes the Mission Creek, the Banning, the Garnet Hill, and the Indio Hills Faults and associated folds (fig. 48). These faults act as barriers to ground-water flow and, combined with constrictions in the basin width and changes in permeability of the water-yielding units, have compartmentalized the Coachella Valley into the Desert Hot Springs, the Mission Creek, the Garnet Hill, and the Whitewater River ground-water subbasins (fig. 48). The Whitewater River subbasin is the largest of the four and contains the most productive aquifer. 
The Gulf of California and its landward extension, the Salton Trough, are structural, as well as topographic, depressions beneath which consolidated rock is thousands to tens of thousands of feet lower than the consolidated rock in the bordering mountains. The Gulf of California and the Salton Trough formed during late Cenozoic time as a result of spreading of the ancient sea floor. 
The aquifer system in the Salton Trough ranges from unconfined in the periphery of the trough to confined in the central part. It consists of basin-fill deposits of Quaternary and Tertiary age (fig. 49); these deposits of alluvium are underlain by rocks of pre-Tertiary age that are referred to as the ³basement complex.² Although the basin fill probably is more than 20,000 feet thick, the water-yielding parts of the basin fill extend only to depths of a few thousand feet. The water at greater depths is too saline for most uses, and the hydraulic connection between the shallow and deep deposits is poor. 
Near the margins of the Imperial Valley, the basin-fill deposits were derived from the adjacent mountains and are mostly coarse sand and gravel. Deposits in the central part of the valley consist mostly of fine-grained sand, silt, and clay that were deposited by the Colorado River. In the eastern and western parts of the Imperial Valley, wells that are open to several hundred feet of the basin-fill deposits yield moderate to large volumes of water. Transmissivity values of 20,000 to 30,000 feet squared per day are characteristic of these deposits, and wells that yield 50 gallons per minute or more per foot of drawdown are attainable. In the central part of the Imperial Valley, transmissivity values of 150 to 1,500 feet squared per day were calculated from two aquifer tests of wells completed in the upper 500 feet of the fine-grained deposits. 
Aquifer material in the Coachella Valley is mostly coarse-grained sediments, and the aquifer is generally unconfined. These deposits are more than 3,000 feet thick, are moderately to highly permeable, and yield large quantities of water to wells. Transmissivity is greatest in the central part of the valley from Palm Springs to the Salton Sea because of the great thickness of permeable deposits. Maximum transmissivity values of about 25,000 feet squared per day are similar to those reported from the Imperial Valley. 
Unlike the other valleys discussed above, the most important source of ground-water recharge to the Imperial and the lower Coachella Valleys is the Colorado River, not runoff from the surrounding mountains. Minor sources of recharge are ground-water inflow from adjacent areas, infiltration of precipitation that falls on the valley floor, and local runoff from the mountains that border the area. 
The Colorado River has been a source of recharge to the aquifer system of the Imperial Valley since the river delta built to a height sufficient to separate the Gulf of California from the Salton Trough. As the delta was built, natural levees beside the river channel kept the Colorado River above the land surface altitude in much of the valley. Under natural conditions, water from the river seeped downward through the river bed and then moved laterally to recharge the aquifer in the Imperial Valley. The water also moved into the Imperial Valley from the Mexicali Valley to the south and through a section of alluvium northeast of the Cargo Muchacho Mountains. 
Since 1901, recharge to the shallow part of the aquifer system under natural conditions has been augmented by percolation of water imported from the Colorado River beginning in 1901. Originally, water was diverted about 10 miles southwest of Yuma, Ariz., at the confluence of the Colorado River and Mexico¹s Alamo Canal, and was delivered to the Imperial Valley through the Alamo and the New Rivers (fig. 46). The completion of the All American Canal (fig. 46), which permitted the diversion of Coloraro River water to the Imperial Valley through a canal located entirely in the United States rather than along a route that passed through the Mexicali Valley, greatly increased the opportunity for ground-water recharge. The All American Canal became the sole means for diverting Colorado River water to the Imperial Valley in February 1942. Six years later, the Coachella Canal was completed and thereafter supplied water to the lower part of Coachella Valley. 
The canals, which are as much as 200 feet wide, are major sources of recharge because they are unlined, flow across many miles of sandy terrain (especially in the eastern part of the Imperial Valley), and are much higher in altitude than the general ground-water levels along their course. The leakage from the canals almost immediately caused ground-water mounds to form beneath the canals, and, over time, ground-water levels rose to the water level in the canals. The leakage also spread horizontally, thereby causing water levels to rise over large areas. Water levels eventually rose to the point that much of the leakage, especially from the All American Canal, was discharged to drains and areas of natural discharge, rather than continuing to add to the quantity of ground water stored in the aquifer system. 
The rise in water levels that resulted from leakage from the easternmost canals between 1939 (before the canals were completed) and 1960 is shown in figure 50. The water-level rise along the All American Canal was generally more than 40 feet, and the rise along the Coachella Canal was about 40 feet near the junction of the canal with the Colorado River and gradually increased northward to more than 70 feet. Throughout most of the length of the East Highline Canal, which began operating in 1942, the original water table was shallow, and the water-level rise was small. 
Water losses along selected reaches of the All American and the Coachella Canals are shown in figure 51. The annual flows are generally 3,000,000 to 4,000,000 acre-feet in the reach of the All American Canal and are about 500,000 acre-feet in the reach of the Coachella Canal. From 1950 through 1967, the average leakage from the two reaches was about 215,000 acre-feet per year. 
The general direction of ground-water movement in the basin-fill aquifer of the Imperial Valley and the lower part of the Coachella Valley is shown by the arrows superimposed on the water-level contour lines in figure 52. The contours were based on water levels in 1965 in wells completed in the main water-yielding zones.
The broad ground-water mound in the southeastern part of the valley is the result of leakage from the All American and the Coachella Canals. Between the canals, the direction of ground-water movement is principally westward, but south of the All American Canal, the movement is toward the Mexican border. Away from the canals, ground water moves generally toward the axis of the valley and then northwestward to the Salton Sea. The principal area of discharge is the central, cultivated part of the valley. Substantial amounts of ground water are discharged to the Alamo River, as indicated by the closely spaced contour lines on the eastern side of the river and the change in direction of the contours which indicates that the ground water flows primarily northward. Ground water also discharges to the New River, but the configuration of the contour lines, which show a relatively wider spacing and moderate upstream displacement, indicates that considerably less ground water moves to the New River than to the Alamo River. 
 For ground water that moves from the adjacent mountains toward the center of the valley, a wide range of contour spacing can be seen in figure 52 that indicates changes in the hydrogeologic conditions within the aquifer. Some of the seemingly abrupt changes in contour spacing are caused by faults that are barriers to ground-water flow. The closely spaced contours west of the Coachella Canal, near Niland, and west of the Salton Sea, near Salton City, result from ground-water flow through deposits with low permeability.
Ground-water movement in and through the basin-fill aquifer of the upper Coachella Valley is affected by the San Andreas Fault System. Faults that are part of this system, as well as constrictions in basin width and changes in permeability of the water-yielding units, have compartmentalized the valley into four ground-water subbasins (fig. 48). 
The general direction of ground-water movement in the upper Coachella Valley was determined from water-level contour maps of the basin for 1936 and 1967 (figs. 53 and 54). Water-level contours during the autumn of 1936 are shown in figure 53. The water-level profile for 1936 (fig. 55) shows the slope of the ground-water surface to be very steep in places and to exceed 50 feet per mile from the San Gorgonio Pass to Windy Point. This steep gradient decreased to less than 10 feet per mile just south of Palm Springs because of the increased width of the basin. From the Thousand Palms area southward, the gradient was about 20 feet per mile. 
A water-level profile for 1967 (fig. 55) shows that while water levels declined in most places in the valley, the steep water-level gradient from the San Gorgonio Pass to Windy Point remained the same as in 1936. However, the water-level gradient at the south boundary decreased to about 10 to 15 feet per mile from the 1936 gradient of 20 feet per mile. The lowered water-level altitudes resulted from increased withdrawals within the Whitewater River subbasin. Water levels for 1967 were about the same as those for 1936 in the extreme southern part of the upper valley as a result of leakage from the Coachella Canal.
The general direction of ground-water flow in 1967 is shown in figure 54 by arrows that represent flow lines, which are the shortest possible paths between adjacent water-level contours. Ground-water movement in the Whitewater River subbasin was primarily parallel to the axis of the valley from Windy Point to Indio. The flow lines near the faults that mark the borders of the subbasins indicate that ground water flowed across the faults; exactly how much is unknown. 
Ground-water development in the Salton Trough has been primarily in the Coachella Valley. The growth of agriculture and, since the early 1950¹s, of tourism has drawn heavily on the ground-water resources of the upper Coachella Valley. Ground-water levels have declined as annual withdrawals increased more than tenfold during 1936 to 1967 (fig. 56). In the lower Coachella Valley, concern over the diminishing ground-water supply as a result of agricultural development prompted the construction of the Coachella Canal. Water delivery began in 1949 when large quantities of Colorado River water were brought to the area between Indio and the Salton Sea. However, the upper Coachella Valley received only small quantities of the canal water and, consequently, water levels continued to decline as ground-water use increased in that area.
Water levels began to decline before 1945 (fig. 57) in the southernmost part of the upper Coachella Valley but did not change significantly in the remainder of the upper valley until about 1945 when major withdrawals began (fig. 56). Water levels continued to decline throughout most of the upper valley through 1965 with the exception of an area near the southern boundary where water levels had ceased to decline and began rising. The rise is documented by observation wells 7 and 8 in 1949 and 1954, respectively (fig. 57). The water-level rise in these two wells can be attributed to water from the Coachella Canal leaking downward to the aquifer. The effect of leakage appeared to be moving up the valley in 1955, as indicated by the hydrograph of well 6 (fig. 57), in which water levels ceased to decline, indicating recharge to the aquifer by leakage. 
From 1936 to 1967, water-level changes in the upper Coachella Valley (fig. 58) were most prominent in the White-water River subbasin. The Palm Springs area in that subbasin showed the largest water-level decline (nearly 80 feet) because of a concentration of withdrawals in an area with a relatively low aquifer storage capacity located near where the aquifer abuts the nearly impermeable bedrock of the San Jacinto Mountains. Decreases in water levels were probably amplified between 1946 and 1967 because of a prolonged drought in the upper Coachella Valley, as in most of southern California, during those years. A representative plot of the departure from average precipitation in the San Jacinto, the Santa Rosa, and the San Bernardino Mountains (fig. 59) indicates that during the dry period from 1947 through 1964 only four years were wet‹1952, 1954, 1957, and 1958. This extended dry period greatly reduced the natural inflow available to the valley. The other three ground-water subbasins showed very little water-level decline because withdrawals in those subbasins were small. A well field for the town of Desert Hot Springs on the east side of the Mission Creek Fault showed some decline because the storage capacity of the aquifer in that area is limited by the juxtaposition of the aquifer with a relatively impermeable fault and the nearly impermeable bedrock of the Little San Bernardino Mountains.
Connected Basin Flow Systems
Although most flow systems are confined to one or two basins in the Basin and Range area, several basins are linked together in an extended ground-water flow system in places (fig. 60). In the majority of the basins, flow passes through the basin-fill sediments that cover the valley floors, as in the Humbolt system. However, where carbonate rocks underlie the basins, data indicate that some basins are hydrologically linked by the carbonate rocks and that large quantities of ground water flow through them and discharge through the overlying basin-fill sediments to large springs. Because few wells are drilled into the carbonate rocks, data are scarce and several assumptions have been made to account for flow in these rocks. One assumption is that the carbonate rocks and the basin-fill deposits form a single hydrologic unit. At locations where wells have been drilled in both rock types, the water levels in each aquifer have been similar. Another piece of evidence that the two rock types act as one hydrologic unit comes from the Ash Meadows area in southern Nevada. Irrigation wells in that area that withdrew water from the basin-fill deposits drew down water levels in the carbonate-rock aquifers more than 1 foot from 1969 to 1972. 
Thick sequences of carbonate rocks underlie most of the alluvial basins within the Basin and Range area in eastern Nevada and southeastern California (fig. 60); these rocks also extend into western Utah, northwestern Arizona, and southeastern Idaho. The carbonate rocks have been faulted, deformed, and eroded through geologic time; original thicknesses of up to 40,000 feet have been reduced by one-half or more. Consequently, most of these rocks are in isolated blocks that form individual aquifers with areal dimensions of only a few square miles. In Nevada, however, the carbonate rocks form a north-south section of aquifer, or ³central corridor² (fig. 61), that is generally laterally continuous for more than 250 miles. The southern part of this corridor has been most studied, and two major flow systems have been identified. In both flow systems, ground water is recharged in east-central Nevada. In one system, ground water discharges at Ash Meadows and Death Valley and, in the other, primarily at Muddy River Springs (fig. 61). 
 These aquifer systems contain the closest approximation to regional flow systems in the Basin and Range area. Flow is transmitted along several valleys through the basin-fill and the carbonate-rock aquifers. However, an insufficient number of wells have been drilled in the carbonate rocks to permit a complete description of the flow system. Discharge from large springs in the alluvial basins is always in areas underlain by carbonate rocks.  
The most complete description of the carbonate-rock flow system is from the ancestral White River/Muddy River Springs area, which consists of 13 valleys in southeastern Nevada (fig. 62). During Pleistocene time, when the climate was more humid than at present, the White River drained six of these valleys and left a wash along its drainageway that is the lowest point anywhere along the drainage system. Five of the remaining seven valleys‹Long, Jakes, Cave, Dry Lake, and Delamar‹are topographically closed and have no streams that flow out of them. The Garden and the Coal Valleys form a topographically closed unit, in which streams in the Garden Valley drain into the Coal Valley. The entire 13-valley drainage area is bounded by mountains that range in altitude from about 7,000 feet in the south to more than 9,000 feet in the north (fig. 62). 
Consolidated rocks (fig. 63) of Paleozoic and Tertiary age form the boundaries of and underlie the basins, which are filled with unconsolidated deposits of Tertiary and Quaternary age. The principal water-yielding rocks of Paleozoic age are limestone and dolomite, which are bounded above and below by confining units of shale, sandstone, and quartzite. The rocks of Tertiary age are primarily tuff and welded tuff and generally transmit little water. However, in areas where the welded tuffs have been fractured, they can yield large quantities of water. The basin-fill deposits consist of coarse sand and gravel at the valley margins and grade to fine silt and clay near the center of the valleys. Valleys that were transversed by the White River during Pleistocene time contain channel deposits of sand and gravel in some places along the ancestral course of the river near the center of the valley. 
Data used to determine the movement of ground water in the White River/Muddy River Springs area are from wells completed in basin-fill sediments and carbonate rocks, altitudes of spring orifices, and several mine shafts in carbonate rocks in the bordering mountains where water levels have been measured. The overall pattern of ground-water movement is shown in figure 64. Ground water moves southward from Long Valley to the White River Valley. The White River Wash defines the principal avenue of flow from the White River Valley to the upper Moapa Valley. Springs issue at several places along the White River Wash wherever the water table is at or near the land surface. Valleys that border the main drainageway of the White River‹the Garden, the Coal, the Cave, the Dry Lake, and the Delamar Valleys‹do not drain directly to the main drainage-way but rather to the south where the ground water leaves them. 
The topography of the drainage area controls the movement of ground water at a regional scale. The flow is in a southerly direction along the axis of the valley as shown by the decrease in water levels along the profile in figure 65, and is generally within tens of feet of the land surface in the center of the valleys. Three areas along this profile show that water levels may be hundreds of feet below land surface. The deep water table in Jakes Valley is caused by the high altitude of the valley floor. Some interpretations of the flow system have identified Jakes Valley as the northern limit of the White River/Muddy River Springs flow system. The deep water table in the Pahroc Valley and at the end of the Pahranagat Valley is thought to be along faults where the fractured rocks have been either cemented to form a barrier to flow or partly dissolved downgradient of the faults so that the permeability is greater and water is more quickly transmitted down-valley from the faults. 
The intervalley ground-water flow pattern and the estimated water budget for the flow system are shown in figure 66. The estimated recharge to the valleys was based on the relation of altitude to annual precipitation (table 2). Because of the orographic effect of the mountains, the amount of precipitation increases as the altitude increases. Conversely, the amount of evapotranspiration decreases as altitude increases, as a consequence of lower temperatures at higher altitudes. Because of higher altitudes and lower temperatures in the northern one-half of the White River/Muddy River Springs drainage area, about 70 percent of the recharge is estimated to be in this one-half of the area. From measurements of spring discharge, about 62 percent of the discharge of the system has been determined to be in the Pahranagat and the upper Moapa Valleys in the southern one-half of the area. This distribution of recharge and discharge is additional evidence of interbasin flow. The substantial amount of ground water that moves from the northern to the southern one-half of the area must flow through the carbonate rocks.
Interbasin flow within the carbonate rocks also is indicated by the water chemistry of perennial springs. Three classifications of the types of flow systems that feed the springs have been defined in the White River/Muddy River Springs area‹regional, large local, and small local. Chemical analyses of water that discharges from perennial springs indicates that increased concentrations of dissolved constituents associated mostly with carbonate rocks are observed with regional systems characterized by long flow paths. Gypsum, anhydrite, halite, and scattered clay minerals in the carbonate rocks are dissolved more readily than carbonate minerals and they contribute sodium, potassium, chloride, and sulfate ions to ground water. Concentrations of these ions increase the longer the ground water is in circulation through the carbonate rocks. At each perennial spring that issues from the carbonate rocks, an indication of the relative distance the water has traveled can be determined by the concentration of these ions. Also, the temperature of the water discharged from a spring indicates the probable depth of ground-water circulation; higher temperatures are associated with the deep circulation of regional systems. 
The concentration of tritium in water that issues from perennial springs has been used to determine places where ground water has had a short residence time in an aquifer. Before 1954, tritium, a hydrogen isotope, was scarce in the atmosphere. Beginning in 1954, detonation of thermonuclear bombs released large concentrations of tritium to the atmosphere in the northern hemisphere. Springs with water that contains tritium in concentrations of 200 units or more (fig. 67) are likely to be discharging a large percentage of ground water that has been in the aquifer for only a short time, perhaps only a few months.
A plot of concentrations of sodium and potassium ions against chloride and sulfate ions in water from the carbonate rocks in Nevada (fig. 68) shows the relation of these ions to the regional, large local, and small local flow systems. Regional flow systems are characterized by interbasin flow, long flow paths, and one or more local systems that feed the regional system. Springs connected to these systems have large perennial discharges and small seasonal ranges in discharge. Discharge waters contain relatively large concentrations of sodium, potassium, chloride, and sulfate ions but have small concentrations of tritium (fig. 68). Some springs discharge thermal water (greater than 80 degrees Fahrenheit). 
Large local flow systems are characterized by predom-inantly interbasin flow and flow paths that are typically confined to one basin. Springs connected to these systems have moderate to large discharges and moderate seasonal ranges in discharge. Discharge waters contain moderate concentrations of sodium plus potassium and chloride plus sulfate (fig. 68) and no significant concentrations of tritium. Discharge waters have temperatures that typically range from 50 to 60 degrees Fahrenheit. 
Small local flow systems are generally characterized by very short flow paths, usually no more than a few miles in length. Springs connected to these systems have highly variable annual ranges in discharge. Discharge waters have small concentrations of dissolved sodium plus potassium and chloride plus sulfate, large concentrations of tritium, and water temperatures that commonly approach average air temperatures.
The evidence of regional flow in the carbonate rocks can be used to evaluate the water-supply potential of the White River/Muddy River Springs Basin. Even if the carbonate-rock aquifer, as yet undeveloped, can supply additional water, the effect of such development upon the aquifer remains uncertain. The effect could be minimal if development can merely capture ground water currently being lost to evapotranspiration in the southern part of the White River/Muddy River Springs area. More information about the aquifer system, particularly data that pertains to aquifer boundaries, is needed in order to accurately determine the potential effects of development

Fresh Ground-Water Withdrawals
An estimated 1,760,000 acre-feet per year of freshwater was withdrawn from aquifers in the Basin and Range area during 1985 (fig. 69). Withdrawals for irrigated agriculture accounted for almost 77 percent of the total. Public supply accounted for almost 18 percent of withdrawals. All other categories totaled less than 6 percent of fresh ground-water withdrawals. 

Ground-Water Quality
The quality of ground water in unconsolidated deposits in the Basin and Range area varies from basin to basin; dissolved-solids concentrations range from less than 500 milligrams per liter (freshwater) to more than 10,000 milligrams per liter, as shown in figure 70. Generally, at the basin margins and on the slopes of alluvial fans, the ground water is fresh. Locally, saline water is present near some thermal springs associated with indurated sedimentary rocks or igneous rocks and in areas where the basin-fill aquifers include large amounts of soluble salts, such as in the upper and middle parts of the Humbolt River Basin. In discharge, or sink, areas such as the Carson and Salton Sinks and Death Valley, however, the dissolved-solids concentration can exceed that of ocean water (about 35,000 milligrams per liter). The ground water beneath playas in small closed valleys may be brackish, but ordinarily the dissolved-solids concentrations are not as large as those in water from the major terminal sinks. Although highly mineralized water is common beneath playas, a deeper freshwater flow system might be present in some areas. For example, water from a well 1,200 feet deep on the northern margin of a playa in a valley near Tonapah, Nev. (fig. 70), has a dissolved-solids concentration of less than 350 milligrams per liter. This concentration apparently reflects deep freshwater circulation in the basin-fill aquifer. In valleys with subsurface discharge into an extensive ground-water system, the water throughout the basin-fill aquifer is generally fresh. 
Introduction
The Central Valley of California (fig. 71) contains the largest basin-fill aquifer system in Segment 1. The valley is in a structural trough about 400 miles long and from 20 to 70 miles wide and extends over more than 20,000 square miles. The trough is filled to great depths by marine and continental sediments, which are the result of millions of years of inundation by the ocean and erosion of the rocks that form the surrounding mountains. Sand and gravel beds in this great thickness of basin-fill material form an important aquifer system. From north to south, the aquifer system is divided into the Sacramento Valley, the Sacramento­San Joaquin Delta, and the San Joaquin Valley subregions, on the basis of different characteristics of surface-water basins.
The Central Valley is one of the most important agricultural areas in the world. No single region of comparable size in the United States produces more fruits, vegetables, and nuts. More than 7 million acres are currently (1995) under irrigation. During 1985, crop irrigation accounted for 96 percent of the surface water and 89 percent of the ground water withdrawn in the Central Valley. 
Discovery of gold in the Sierra Nevada and the subsequent proliferation of hydraulic mining operations provided the impetus for the construction of a surface-water diversion system that consisted of hundreds of miles of canals used to transport water to where it was needed for gold-washing operations. This was the beginning of the valley¹s modern-day aqueduct system, which has become vital to the agricultural economy.
Fertile soil, favorable climate, abundant water, and rapid population growth in the Central Valley encouraged the development of agriculture, which soon became one of the major industries of California. Surface water satisfied most irrigation needs until the late 19th century, when a rapid increase in irrigated acreage produced a demand for water that exceeded the surface-water supply, and ground-water supplementation became necessary; the drought of 1880 was a major stimulus for ground-water development. Wells were used to supplement less dependable surface-water supplies and to provide water where surface-water diversion canals had not been constructed. Shallow ground water was obtained easily in 1880, and artesian pressure was sufficient to produce flowing wells in much of the valley. After 1900, ground water gradually became a more significant part of the total irrigation supply and, eventually, the large number of wells reduced artesian pressure to such an extent that it became necessary to install pumps in order to obtain water. 
The invention of the deep-well turbine pump around 1930 allowed withdrawals from greater depths, which encouraged further development of ground-water resources for irrigation. Withdrawals increased sharply during the 1940¹s and 1950¹s, and averaged about 11.5 million acre-feet per year by the 1960¹s and 1970¹s, which was approximately 20 percent of the total irrigation withdrawals for the United States at that time. Withdrawals reached a maximum of 15 million acre-feet per year during 1977, a drought year. During the 1960¹s and 1970¹s, withdrawals greatly exceeded recharge, and water levels declined precipitously, as much as 400 feet in places. The declines caused a major reduction in the amount of ground-water in storage and resulted in widespread land subsidence, mainly in the western and southern parts of the San Joaquin Valley. 
Increased rainfall and construction of additional surface-water delivery systems halted most of the serious water-level declines after 1977, and water levels recovered to pre-1960 levels. The network of aqueducts in the Central Valley is currently (1995) sufficient to provide one-half or more of the water needed for irrigation in years of average or above-average precipitation. In dry years, however, reliance on ground-water supplies is greater, and aquifers might again be subject to withdrawals in excess of recharge during a severe drought, such as that of 1976­77. 
The Central Valley is bounded on the west by the Coast Ranges and on the east by the Cascade Range and the Sierra Nevada. The valley has only one surface-water outlet, the Carquinez Strait east of San Francisco Bay. Much of the valley is surrounded by dissected uplands formed by erosion of coalesced alluvial fans at the base of the mountains (fig. 72) where the terrain ranges from hilly to slightly rolling. The valley floor, which consists primarily of alluvial deposits and flood-plain deposits of the major rivers, is relatively flat to gently rolling and is generally below an altitude of 500 feet. Lake beds in the southern end of the valley become partially to completely flooded in wet years. A prominent feature, Sutter Buttes, which is the remnant of a volcanic plug, rises nearly 1,500 feet above the valley floor in the central Sacramento Valley. 
The Sacramento River drains the northern end of the Central Valley, and the San Joaquin River drains much of the middle third. The two rivers join in the Sacramento­San Joaquin Delta and empty into the upper end of San Francisco Bay. The southern end of the valley is occupied by the Tulare Basin, in which drainage is completely internal and the inflowing water is removed by evapotranspiration.
The climate of the Central Valley is Mediterranean and Steppe, characterized by hot summers and mild winters, thus allowing for a year-around growing season; at least one crop is under cultivation at all times. About 85 percent of the precipitation falls from November to April. Most of the precipitation that falls on the valley floor evaporates before it can infiltrate downward to become recharge. Much of the moisture that moves inland from the Pacific Ocean is intercepted by the Coast Ranges, so that annual precipitation in the valley is relatively low. Annual precipitation decreases from north to south, with an average of about 23 inches in the northern part of the Sacramento Valley, to about 6 inches in the southern part of the San Joaquin Valley. Rainfall amounts vary greatly from year to year. Annual precipitation is exceeded by potential evapotranspiration throughout the entire valley, which causes a net annual moisture deficit. 
In contrast, the mountains that surround the Central Valley intercept moisture from eastward-moving weather systems and have an annual surplus of moisture in the form of rain and snow. Precipitation can exceed 80 inches annually in the Sierra Nevada. Annual runoff from rainfall and snowmelt is approximately 32 million acre-feet; most of the runoff originates in the Cascade Range and the northern Sierra Nevada (fig. 73). This water flows to the valley in perennial streams and provides nearly all the average annual 12 inches of recharge the valley aquifer system receives. Runoff from the Coast Ranges is principally on the western slopes to the Pacific Ocean.
Geologic Setting
The Central Valley and surrounding area is the product of a complex series of geologic events. The surrounding area has undergone mountain building, faulting, and erosion, and the valley has been inundated several times by the Pacific Ocean. 
The Sierra Nevada, which forms the eastern side of the valley, is the eroded edge of a huge tilted block of crystalline rock that also partially defines the base of the valley sediments (fig. 74). Embedded in the granite and related plutonic rocks of the mountains are metamorphosed sedimentary and volcanic rocks of Ordovician to Late Jurassic age. The uplift that formed the Sierra Nevada probably took place between Late Jurassic and Late Cretaceous time. 
The northeast corner of the basin is the southern terminus of the Cascade Range. This is an area of lava plateaus and volcanos, some of which have been active in modern times. Geologically, this area of the basin is relatively young; most of the volcanic activity was during late Tertiary to Holocene time. 
The northwest corner of the Central Valley is bounded by metamorphosed volcanic rocks of Paleozoic and Mesozoic age. These rocks form a minor part of the valley boundary. The western side of the valley is bounded by the Coast Ranges, which are formed primarily of folded and faulted marine sedimentary rocks of Mesozoic age, which were uplifted during Tertiary time. Mesozoic marine sedimentary rocks and continental deposits underlie the Tehachapi Mountains that bound the valley to the south.
A huge volume of sediments, which is as thick as about 50,000 feet in the Sacramento Valley and about 32,000 feet in the Tulare Basin, fills the Central Valley. These sediments are marine and continental in origin; the marine sediments are the product of deposition during inundations by the Pacific Ocean, and the continental sediments were derived by erosion of the rocks that formed the surrounding mountains. 
The ancestral Central Valley was, at least in part, inundated by the Pacific Ocean until 2 million to 3 million years ago. The location, depth, and age of marine sediments in the valley indicates that nearly the entire valley was covered by the sea during Paleocene and Eocene time (fig. 75). As sea level declined, the area covered by the ocean decreased until only the southern end of the basin was still under water in Pliocene time. During Pleistocene and Holocene time, the sea completed its retreat, and all oceanic deposition ceased. In total, the ocean left behind deposits that ranged in thickness from about 25,000 feet in the Sacramento Valley to about 20,000 feet in the San Joaquin Valley. These deposits are mostly consolidated and have minimal permeability. 
From the time when the valley first began to form, sediments derived from erosion of igneous and metamorphic rocks and consolidated marine sediments in the surrounding mountains have been transported into the valley by streams. These continental sediments are as thick as 9,000 feet at the southern end of the valley and have an average thickness of about 2,400 feet (fig. 76). The continental sediments consist mostly of fluvial, basin-fill, or lake deposits of sand and gravel interbedded and admixed with clay and silt (fig. 77). Depending upon location, deposits of fine-grained materials‹mostly clay and silt‹make up as much as 50 percent of the thickness of the valley-fill sediments.

Geohydrologic Setting
Drainage Basins
Three hydrologic subregions coincide with drainage basins within the Central Valley (fig. 71). These subregions are hydraulically connected and compose the Central Valley aquifer system and associated surface-water drainages. The northernmost subregion is the Sacramento Valley, which extends over the northern one-third of the Central Valley and is drained by the Sacramento River. Although the Redding Basin extends over about 500 square miles at the northern end of the Sacramento Valley and is a topographic extension of the valley, it is not included as part of the Central Valley aquifer system because of its separate ground-water flow system. Adjoining the Sacramento Valley to the south is the Sacramento­San Joaquin Delta subregion, where a network of meandering channels has formed at the junction of the Sacramento and the San Joaquin Rivers. The southernmost subregion is the San Joaquin Valley, which extends over two-thirds of the Central Valley. The San Joaquin River drains the northern part of the San Joaquin subregion; the southern part, which is called the Tulare Basin, is characterized by interior drainage. The Tulare Basin is named for Tulare Lake, a lake that covered much of the basin during the Pleistocene Epoch.
Under natural, or predevelopment, conditions, recharge from rainfall and snowmelt entered the aquifer system as seepage from streams that channel runoff from the surrounding mountains into the valley. Most recharge is at the margins of the valley, and the ground water moves in the subsurface to lower altitudes and discharges into surface-water bodies that drain each basin. 
Aquifers and Confining Units 
The consolidated volcanic and metamorphic rocks that surround and underlie the Central Valley are almost impermeable, and flow through them is not significant. Little water flows through the extensive deposits of consolidated marine and mixed marine and continental sediments that overlie the crystalline rocks (fig. 77) because the permeability of the deposits is generally minimal. The marine sediments usually contain saltwater or brine, but near the northwestern, western, and southeastern margins of the San Joaquin Valley, some freshwater is withdrawn from these deposits. 
The Central Valley aquifer system is formed primarily of sand and gravel with significant amounts of silt and clay, all of which have been eroded mainly from older rocks at the boundaries of the valley. The environments in which the continental sediments were deposited varied, but most were deposited in fluvial environments; however, the deposits contain some lacustrine beds. Locally, volcanic rocks and dune deposits are part of the aquifer system. Specific geologic formations can be related to specific aquifers within the Central Valley aquifer system only with difficulty because many of the formations are lithologically similar, and cannot be distinguished easily in the subsurface. 
Beds and lenses of fine-grained materials, such as silt and clay, constitute a significant percentage of the Central Valley aquifer system. In most parts of the valley, fine-grained materials compose 50 percent or more of the aquifer system. The most extensive clay bed, which is informally named the ³E-clay² (fig. 77), consists primarily of the Corcoran Clay Member of the Tulare Formation and underlies much of the western San Joaquin Valley. Because beds of silt and clay do not readily transmit water under natural conditions, they act as barriers to vertical flow and cause differences in hydraulic head with depth. 
Early investigators thought that the Sacramento Valley contained a single unconfined aquifer and that the San Joaquin Valley contained an upper unconfined to semiconfined aquifer separated from a lower aquifer confined by the Corcoran Clay or ³E-clay² (fig. 78). However, recent investigations indicate that the Central Valley contains a single heterogeneous aquifer system that contains water under unconfined, or water-table, conditions in the upper few hundred feet; these conditions grade into confined conditions with depth. The confinement is the result of numerous overlapping lens-shaped clay beds. Geophysical well logs indicate that the ³E-clay,² although probably the largest single confining bed, constitutes only a small percentage of the total thickness of clay layers in the aquifer system. This indicates that the significance of the ³E-clay² as a barrier to vertical flow may have been exaggerated. Further, the difference in hydraulic head directly above and below the ³E-clay² is small when compared to head differences within intervals of the deep parts of the aquifer system. 
Ground-Water Flow System
Before development began, the aquifer system was under steady-state conditions in which natural recharge balanced natural discharge. Ground water in the shallow part of the aquifer system flowed from areas of high altitude at the valley margins, where most of the recharge took place, downgradient to discharge into rivers and marshes near the valley axis (fig. 79). The aquifer system was recharged primarily by streams emanating from the Coast and Cascade Ranges and the Sierra Nevada. Most of the recharge was in the northern and eastern parts of the valley. Precipitation falling on the valley floor during the rainy season provided only a small part of the total recharge. Ground water that was not evaporated or transpired by plants discharged either into the Sacramento and the San Joaquin Rivers that drained to San Francisco Bay or into the Tulare Basin from which it was eventually removed by evaporation or transpiration. The areas of recharge and discharge in the Central Valley before development are shown in figure 80. 
Under predevelopment conditions, the hydraulic head in the shallow water-table aquifer where water entered the aquifer system at the valley margins was greater than the head in the deeper confined aquifer; thus, ground water moved downward (fig. 81). Conversely, the head gradient was reversed where water left the aquifer, typically by discharge to surface-water bodies, and the hydraulic head in the water-table aquifer was less than that in the confined aquifer. The difference in hydraulic head created upward movement of the ground water toward rivers and marshes (fig. 81). Precipitation that fell on the valley floor and was not lost to evapotranspiration recharged the water-table aquifer and moved down the head gradient toward the rivers and surrounding marshes. 
Upward vertical flow to discharge areas from the deep confined aquifer was impeded by confining clay beds, which caused a pressure head in the deep parts of the aquifer system. Because of the pressure head, wells that penetrated the deep aquifer in low-lying areas near the rivers and marshes flowed during the early years of ground-water development in the valley. 

Postdevelopment Ground-Water Flow System
By the early 1960¹s, intensive ground-water development had significantly lowered water levels and altered ground-water flow patterns in the Central Valley aquifer system. By far the most dramatic impact of development was in the San Joaquin Valley, where water-level declines in the confined part of the aquifer system were locally more than 400 feet (fig. 82). Although predevelopment flow was toward the San Joaquin River throughout most of the basin, large withdrawals from deep wells in the western and southern parts of the aquifer system changed the direction of horizontal flow in the confined part of the system until the water moved toward the withdrawal centers (fig. 83). Also, because the magnitude of the withdrawals caused hydraulic heads in the confined parts of the aquifer system to fall far below the altitude of the water table (fig. 84), the vertical hydraulic gradient was reversed over much of the San Joaquin Valley. As a result, much of the water in the upper unconfined zone of the aquifer system that flowed laterally toward the river under predevelopment conditions leaked downward through the confining beds into the lower confined aquifer after development (fig. 85). 
Concurrent with an increase in surface-water imports in the early 1970¹s, ground-water withdrawals in the northern part of the Central Valley aquifer system decreased, which allowed ground-water levels in many areas to recover in the confined part of the aquifer system (fig. 86), in some cases to pre-1960 levels. However, in the San Joaquin Valley large withdrawals continued, especially in the western and southern parts of the valley, and water levels continued to decline. With few exceptions, the ground-water flow patterns in the aquifer system today (1995) are the same as those in the mid-1970¹s.
Ground-water development in the San Joaquin Valley has reduced the effectiveness of the confining beds within the aquifer. Thousands of wells with casings perforated for much of their length have been drilled through the clay confining units. Where these wells are open to the unconfined and confined aquifers, they allow virtually unrestricted vertical flow through the well bore (fig. 87). The amount of water that flows downward through one large-diameter well has been estimated to be equivalent to the natural leakage through the ³E-clay² over an area of approximately 7 square miles. During the peak of the withdrawal season, the net downward flow may be, on average, as much as 0.3 cubic foot per second per well. 

Well Depths and Yields
Well depths in the Central Valley aquifer system are determined by the depth of permeable aquifer material and by the quality of the ground water. In general, wells are usually less than 500 feet deep in the Sacramento Valley but are as deep as 3,500 feet in the San Joaquin Valley. The greater depth of wells is a result of the low permeability of the sands in the unconfined aquifer in the western and southern San Joaquin Valley and of highly mineralized water and water high in selenium in the upper parts of the aquifer system in the western San Joaquin Valley. 
Well yields of more than 1,000 gallons per minute are commonly obtainable throughout the aquifer system. The average yield of wells in the Sacramento Valley is approximately 800 gallons per minute, but yields as large as 4,000 gallons per minute have been recorded. The average yield of wells in the San Joaquin Valley is about 1,100 gallons per minute, and the maximum expected yield is about 3,200 gallons per minute. 

Water Budget of the 
AQUIFER SYSTEM
A water budget is a method of quantitatively accounting for water movement in a hydrologic system. A computer-simulated approximation of the annual Central Valley aquifer system water budget under predevelopment and development conditions is shown in figure 88. The figure depicts only water that circulates through the aquifer system and does not account for water that enters the valley but does not interact with the aquifer system. This excludes most surface-water flow and water that is lost to evaporation almost immediately after it falls on the valley floor. 
Before development, the net circulation through the aquifer system was approximately 2 million acre-feet per year (fig. 88A). Of an average annual 12.4 million acre-feet of precipitation that fell on the valley floor, 10.9 million acre-feet was lost to evaporation because of the arid conditions that characterize the valley; thus, only 1.5 million acre-feet of precipita-tion entered the aquifer system as recharge. Water that moved from surface-water bodies to the aquifer system provided the remaining 500,000 acre-feet per year of recharge. The recharge was balanced by discharge from the aquifer system to rivers (300,000 acre-feet) and evapotranspiration (1.7 million acre-feet). 
Development added two components to the water budget‹withdrawals and return flow from irrigation‹and increased the volume of water that flowed through the ground-water system approximately sixfold from 1961 through 1977. Ground-water withdrawals for irrigation, municipal supply, and industrial use totaled about 11.5 million acre-feet annually. Seepage from irrigation returned about 9 million acre-feet to the ground-water system (fig. 88B). During the period 1961 through 1977, the rate of ground-water withdrawals from the aquifer system was greater than the net recharge from all sources. Withdrawals in excess of recharge resulted in a loss of water from storage in the aquifer of 800,000 acre-feet per year. In the case of the Central Valley aquifer system, some of the loss from storage is permanent because some of the water was removed from beds of fine-grained materials, which, when drained, become compacted and cannot store water again. Compaction of fine-grained materials led to land subsidence in the Central Valley. By the late 1970¹s, however, sufficient surface-water supplies were imported by aqueducts to reduce substantially the volume of ground water that was withdrawn. Although some additional surface water has been imported since 1977 and ground-water withdrawals have slightly decreased, the water budget shown in figure 88B is representative of current (1995) conditions. 
Fresh Ground-Water Withdrawals
Ground-water withdrawal from the Central Valley aquifer system varies seasonally. The highest demand is generally during the peak growing season in spring and summer, which are the driest seasons of the year. Demand for ground water is greatest in the semiarid San Joaquin Valley where natural recharge is least. Withdrawal rates increase significantly during dry years (fig. 89). 
Ground water accounted for only a small part of the water withdrawn for irrigation before 1900. Streams and distribution canals supplied most of the demand. However, the need to con-tinue irrigation in dry years when surface-water supplies are undependable, as well as the expansion of agriculture into areas distant from surface-water sources, prompted increased ground-water development. By the 1960¹s, ground-water withdrawals from the Central Valley aquifer system averaged 11.5 million acre-feet per year, which was one-half or more of the water withdrawn from all sources (fig. 90) and was about 20 percent of the total irrigation withdrawals for the entire United States. During that same period, withdrawals for domestic and industrial uses accounted for about 5 percent of all ground-water withdrawals in the Central Valley. Historically, the largest withdrawal of ground water‹15 million acre-feet‹was during 1977, a drought year. 
Increased importation of surface water for irrigation since 1977, as well as generally wetter weather through 1985, resulted in decreased ground-water withdrawals. During 1985, ground water accounted for only about 37 percent of the total withdrawals in the Central Valley (fig. 91); total ground-water withdrawals were 10.1 million acre-feet (fig. 92). Of that amount, agricultural withdrawal accounted for 8.8 million acre-feet, all of which was used for irrigation (fig. 92A). This amount was about 11.5 percent of all ground water withdrawn in the United States for all purposes during 1985. The remaining 1.3 million acre-feet was used for public supply and industrial purposes, and by domestic and commercial users. Almost one-half of the water was withdrawn from the Tulare Basin (fig. 92B).
Although gains have been achieved by importing surface water from areas of surplus to areas of deficit, projected water needs in the San Joaquin Valley may require temporary withdrawal of ground water in excess of recharge in the future. The California Department of Water Resources has estimated that by 2010 demand for water in the Sacramento Valley, the San Joaquin Valley (exclusive of the Tulare Basin), and the Tulare Basin would be 7, 8, and 9 percent greater, respectively, than 1980 demands. The Sacramento Valley is expected to have sufficient supplies to meet agricultural demand until at least 2010. However, without increased surface-water imports, the San Joaquin Valley (exclusive of the Tulare Basin) and the Tulare Basin might require withdrawals of 150,000 and 2,400,000 acre-feet per year, respectively, in excess of recharge. Those estimates probably underestimate additional increased demand that would result from sustained dry weather. Occasional large withdrawals from an aquifer are a viable solution to the problem of reduced surface-water supplies in dry periods, provided the aquifer is replenished during wet years. However, continual withdrawal of ground water in excess of recharge can increase the cost of pumping, reduce water availability, and, in certain hydrogeologic settings, can cause land subsidence.

Land Subsidence
Land subsidence is widespread in the Central Valley (fig. 93), and has resulted in damage to buildings, aqueducts, well casings, bridges, and highways; has caused flooding; and has cost millions of dollars. The three processes that caused most of the subsidence are oxidation and compaction of peat, hydrocompaction, and compaction of fine-grained sediments due to withdrawal of ground water in excess of recharge. Of these, the third process has caused the most widespread and severe subsidence. 
Human-induced subsidence probably began in the middle 1800¹s when peat soils in marshes of the Sacramento­San Joaquin Delta were first drained for cultivation. In the delta, shallow ground water is drained into ditches to dry the fields before planting and then pumped from the ditches into nearby natural channels. During the growing season, water is siphoned back into the drainage ditches to raise the water table to the root zone. When the peat soils are drained and exposed to the atmosphere they oxidize and compact, and (or) are reduced in thickness by wind erosion; thus, the land surface is permanently lowered with each yearly cycle. To dry the fields each year, the water table must be lowered below that of the previous year, which requires an increase in withdrawals and a decrease in the volume of ground water in storage. Subsidence due to oxidation and compaction of peat soils has lowered the land surface in the delta as much as 6 to 15 feet. 
Hydrocompaction is caused when formerly unsaturated soils become saturated, which allows the soil particles to reorient into a more compact form. Irrigation of clayey alluvial-fan soils has resulted in hydrocompaction and subsidence of 3 to 15 feet on the western and southern margins of the San Joaquin Valley (fig. 93). Soils in many areas crossed by the California Aqueduct were intentionally hydrocompacted before aqueduct construction to avoid subsidence problems. Subsequent subsidence due to hydrocompaction in these areas has been minimal. 
The primary cause of land subsidence in the Sacramento and the San Joaquin Valleys has been the compaction of fine-grained sediments (predominantly clay) in the aquifer system following severe, long-term withdrawal of ground water in excess of recharge (fig. 93). The amount of such subsidence in an area is related to the amount of withdrawal and the percentage of the withdrawal zone composed of clay beds. Compaction occurs when the hydraulic head in the confined parts of the aquifer system is lowered, thus reducing the hydraulic head in the clay beds, which, in turn, reduces the pore pressure in the clay. The weight of overlying sediments compacts the clay and squeezes water out of the clay until equilibrium is reached with the pore pressure in the clay. Compaction seems to happen more readily when the wells are open only to the confined part of the aquifer system than when they are open to the shallow water-table aquifer as well. When ground water is withdrawn from above and below the confining units, head differential is less between the shallow and deep aquifers and reduction in pore pressure in the clay is less.
Subsidence due to compaction of fine-grained sediments began in the San Joaquin Valley in the 1920¹s and in the Sacramento Valley in the 1950¹s. The area most affected has been in the southern and western parts of the San Joaquin Valley (fig. 93). Approximately one-half of the valley, or about 5,200 square miles, had subsided at least 1 foot by 1977; the total volume of subsidence was greater than 17 million acre-feet. The land surface declined nearly 30 feet from the 1920¹s to the late 1970¹s in an area southwest of Mendota (fig. 94). Importation of surface water and reduction in ground-water withdrawals during the 1970¹s slowed or stopped the decline of ground-water levels. In many cases, this allowed recovery to pre-1960¹s water levels and prevented further land subsidence. 
Compaction and