Scientific Investigations Report 2006-5122
U.S. GEOLOGICAL SURVEY
Scientific Investigations Report 2006-5122
Numerous and complex hydrogeologic and hydrologic data used to develop the conceptual model are described in six sections of this report: (1) “Hydrogeologic Framework and Hydraulic Properties,” which defines three hydrogeologic units, and the structural features and hydraulic properties of these hydrogeologic units that affect ground-water flow in the model area; (2) “Model Boundaries,” which defines the location of the three physical and three artificial hydrologic boundaries of the model; (3) “Inflows, Outflows, and Fluxes Across Model Boundaries,” which characterizes the model boundaries and describes inflows, outflows, and fluxes across the boundaries; (4) “Ground-Water Budget,” which summarizes model inflows and outflows and analyzes uncertainties of the water-budget components; (5) “Ground-Water Movement,” which describes the hydraulic gradient, ground-water flow directions, and average linear ground-water velocities in the model area; and (6) “Model Representation of Features Affecting Ground-Water Flow,” which describes the stratigraphic, structural, and hydrologic controls on ground-water flow in the model area. The report concludes with a brief section titled “Implications for Contaminant Transport” which summarizes those features of the conceptual model that most affect interpretations of contaminant transport in the aquifer at the INL and vicinity.
The hydrogeologic framework is primarily based on broad differences in the lithology and large variations in the hydraulic properties of the heterogeneous, anisotropic basalt-flow groups, sedimentary interbeds, and other rocks that allow these rocks to be grouped into three homogeneous and anisotropic hydrogeologic units. In the unsaturated zone and aquifer beneath the INL, basalt makes up about 85 percent of the volume of rocks and sediment makes up most of the remainder (Anderson and Liszewski, 1997, p. 11). These and other rocks form at least 178 basalt-flow groups, 6 andesite-flow groups, 103 sedimentary interbeds, and 4 rhyolite domes (Anderson and Liszewski, 1997, p. 21, table 4). The aquifer includes at least 65 basalt-flow groups, 5 andesite-flow groups, 61 sedimentary interbeds, and 3 rhyolite domes. Detailed surficial mapping and correlation of subsurface stratigraphic units among numerous outcrops and 333 wells at and near the INL (Kuntz and others, 1994; Anderson and others, 1996a) indicated that many of these stratigraphic units are continuous across large parts of the model area. Stratigraphic units were combined by Anderson and Liszewski (1997, p. 14), based on their similar age, into 14 composite stratigraphic units, each made up of from 5 to 90 (Anderson and Liszewski, 1997, p. 14 and table 4) stratigraphic units of similar age (fig. 6).
Structural features in the model area include (1) rhyolite domes (Kuntz and others, 1994), (2) sedimentary troughs (Gianniny and others, 1997), (3) areas of subsidence, uplift, and dipping beds (Anderson and others, 1997, Anderson and Liszewski, 1997), (4) volcanic rift zones, which are broad belts of focused volcanism that generally trend northwestward (fig. 7) and are perpendicular to the direction of regional ground-water flow (Kuntz and others, 1992), and (5) vent corridors, which are narrow zones in and near volcanic rift zones that contain known or inferred volcanic vents, dikes, and fissures (fig. 8) (Anderson and others, 1999, p. 13).
Hydraulic properties of the three hydrogeologic units of the conceptual model reflect the distribution of major rock types, local stratigraphic units, and structural features. Areas of high well density at the INTEC, RWMC, TRA, and TAN allowed evaluation of local variations in hydraulic properties related to the complex geology. In these areas, large variations in hydraulic properties across distances of hundreds to thousands of feet were measured. For example, horizontal hydraulic conductivity in a single vent corridor at the INTEC and TRA varies as much as three to five orders of magnitude across distances of 500 to 1,000 ft (Anderson and others, 1999, p. 27, table 2). These large variations indicate the potential complexity of the aquifer at the scale of an individual INL facility. Although this small-scale complexity cannot be duplicated at the scale of the conceptual model, it can be used as a guide for estimating hydraulic properties in the model with greater precision than was possible for the RASA study.
Hydraulic properties for the eastern SRP aquifer used in the larger scale RASA study were based on specific capacity or transmissivity and storativity estimates from medium- to high-capacity (greater than 50 gal/min) irrigation and industrial supply wells (Mundorff, 1964; Whitehead, 1992, tables 4 and 5, Garabedian, 1992, tables 3 and 4). Transmissivity, horizontal and vertical hydraulic conductivity, storativity, and porosity at the INL and vicinity were estimated at various scales. Core-scale estimates of hydraulic conductivity and porosity are given in Knutson and others (1990, 1992). Transmissivity estimates for single-well aquifer tests (generally less than 40 gal/min) are given in Ackerman (1991, table 2) and Bartholomay and others (1997, table 3). Additional single-well aquifer tests, slug tests, and packer tests are summarized in Welhan and Wylie (1997) and Welhan and others (2002a). Estimates of horizontal and vertical hydraulic conductivity and of storativity from multi-well aquifer tests are given in Spinazola (1994), Frederick and Johnson (1996), and Wood and Norrell (1996).
Hydraulic conductivity of the basalts of the eastern SRP aquifer at the INL and vicinity range from about 0.01 to 24,000 ft/d (Anderson and others, 1999, fig. 9, table 2), more than six orders of magnitude. Almost two-thirds of these estimates exceeded 100 ft/d. Hydraulic conductivities were estimated by dividing transmissivity by the total lengths of open, perforated, or screened intervals. Transmissivities of the eastern SRP aquifer range from 1.1 to 760,000 ft2/d, nearly six orders of magnitude (Ackerman, 1991, table 3). More than 60 percent of these estimates are greater than 20,000 ft2/d (Ackerman, 1991, fig. 10). Garabedian (1992, tables 19 and 20) used a model-calibrated range of 0.45 to 9,500 ft/d for the hydraulic conductivity of all types of basalt in the RASA study.
Estimates of the fractured basalt porosity in the eastern SRP aquifer range from 0.05 to 0.27 (Nace and others, 1959, p. 58-61, table 9; Barraclough and others, 1967, p. 61 and 63; Robertson and others, 1974, p. 176; Robertson, 1974, p. 13, 1977, p. 44-45; Garabedian, 1992, p. 44-46; Ackerman, 1995, p. 10 and 22; Bishop, 1991, p. 77; Knutson and others, 1992, p. 4-21). Estimates of porosity varied greatly because they were dependent on methods, scales, and locations used to determine them. For example, porosity estimated from laboratory measurements of cores from boreholes and outcrops at and near INL facilities were different from estimates derived from larger-scale aquifer model simulations. Because porosity in the eastern SRP basalts derives mainly from interflow zones and their associated rubble, fractures, joints, and vesicles (Hughes and others, 1999, fig. 12), the larger-scale estimates of porosity for these basalts are probably more appropriate for the conceptual model.
Interflow zones may be the most important hydraulic features of the eastern SRP aquifer because of their high permeability and their ability to localize large ground-water fluxes (Welhan and others, 2002a, p. 226). Five interflow zones identified near INTEC, in the upper 200 ft of the aquifer, range in thickness from less than 1-ft thick to more than 18-ft thick (Jones, 1961, p. 36), and the effective thickness of interflow zones may be between 1 and 8.2 ft (Welhan and others (2002b, p. 140). Welhan and others (2002a, p. 226-227; 2002b, p. 137-139) described two types of interflow zones. Type-I interflow zones are hosted in rubble at horizontal contacts between basalt flows and Type-II interflow zones are hosted in fracture networks along the edges of inflating lava flows (Welhan and others, 2002b, p. 136). The largest concentration of basalt fractures in eastern SRP basalts occurs at interflow zones (Welhan and others, 2002a and 2002b). Fractures in Type-I interflow zones consist of tension fractures and fractured flow tops in the upper vesicular zone. Fracturing in Type-II flow zones occur as interconnected networks of fissures along the intensely fractured margins of basalt flows that can remain partially open after burial by younger lava (Welhan and others, 2002a, p. 229).
The hydrogeologic units in the model area are (1) hydrogeologic unit 1, younger rocks of thin, densely fractured basalt (440,000 to 650,000 years old) and interbedded sediment, (2) hydrogeologic unit 2, younger rocks of massive, less densely fractured basalt (650,000 to 800,000 years old) and interbedded sediment, and (3) hydrogeologic unit 3, intermediate-age rocks of slightly altered, fractured basalt (800,000 to 1,800,000 years old) and interbedded sediment (figs. 9 and 10). Other rocks of hydrogeologic importance in the model area are (1) older rocks of intensely altered basalt (older than 1,800,000 years), the top of which is interpreted to form the base of the aquifer, (2) rhyolite domes that penetrate the younger and intermediate-age rocks, and (3) sediment that is distributed throughout the younger, intermediate-age, and older rocks.
Hydrogeologic units 1, 2, and 3 correlate with the regional stratigraphy and the stratigraphy beneath the INL as defined by Anderson and Liszewski (1997) as follows: (1) hydrogeologic unit 1 is equivalent to the Snake River Group and composite stratigraphic units 4 to 6; (2) hydrogeologic unit 2 is equivalent to the Snake River Group and composite stratigraphic unit 7; and (3) hydrogeologic unit 3 is equivalent to the Bruneau Formation and composite stratigraphic units 8 to 14 (fig. 11). Equivalent groups and formations were determined based on broad correlations between cores at the INL and outcrops in southern Idaho (Whitehead, 1992; Anderson and Liszewski, 1997). In the model area these composite stratigraphic units can be saturated (aquifer) or unsaturated (fig. 11).
Younger rocks form the uppermost part of the aquifer in much of the model area and, based on stratigraphic interpretations, intermediate-age rocks make up the largest volume of the aquifer in the model area. Interpreted distribution of older rocks indicates large changes in saturated thickness across the model area that may be related to differential subsidence and uplift (Anderson and Liszewski, 1997, fig. 6). In the southwestern part of the model area the hydrogeologic framework is interpreted as having a zone of differential subsidence and uplift that affects the dip of hydrogeologic units 1 and 2 (fig. 10). This interpretation, based on the trend of composite stratigraphic units 4 through 7 (Anderson and Liszewski, 1997, figs. 21, 23, 25, and 27), is uncertain due to the few core and stratigraphic data available in the southwestern part of the model area.
Hydraulic conductivities for hydrogeologic units 1, 2, and 3 and older rocks were estimated from aquifer tests in wells using straddle packers or having perforated or open intervals only in the respective hydrogeologic unit or older rocks, and hydraulic conductivities of sediment were estimated from laboratory determinations of the saturated hydraulic conductivity of core samples (table 2). Estimated hydraulic conductivity from core-scale tests were not used to estimate the hydraulic conductivities of the hydrogeologic units or older rocks because they reflect the character of the rock matrix and are not representative of the magnitude and spatial variability of hydraulic conductivity estimated from field-scale tests arising from larger-scale heterogeneities (Welhan and others, 2002a, p. 229). These core-scale tests have minimum and median values 3 to 4 orders of magnitude less than field-scale test results (Welhan and others, 2002a, fig. 3).
Younger rocks (hydrogeologic units 1 and 2) make up the unsaturated zone throughout most of the model area and the uppermost part of the aquifer in the central part of the model area (figs. 9 and 10). Sediment content is largest in the younger rocks in a sedimentary trough, known as the Big Lost Trough, in the northern part of the model area. Petrographic analyses of cores of the younger rocks from wells at and immediately north of the INTEC (Lanphere and others, 1993, p. 19-33) indicated that alteration was rare and weak and that secondary mineralization consisted of thin coats of carbonate minerals. The younger rocks also include four rhyolite domes [fig. 9, rhyolite domes A, B, C, and an unnamed dome (not shown in fig. 9) between B and C].
Hydrogeologic unit 1 comprises 17 basalt-flow groups, 4 andesite-flow groups, and 25 sedimentary interbeds (fig. 11) (Anderson and Liszewski, 1997, p. 16-17, figs. 9-15). These thin, densely fractured basalt flows are mostly 14- to 24-ft thick (table 3) tube-fed pahoehoe flows having thin, massive interiors (Anderson and others, 1999). Mean thickness of the basalt flows in hydrogeologic unit 1 is 20 ft.
Hydraulic conductivity and porosity of hydrogeologic unit 1 primarily is controlled by void spaces of interconnected interflow zones. Other factors that probably modify the hydraulic conductivity and porosity of the fractured basalt include near-vent deposits, dikes, fissures, and tension cracks within volcanic rift zones and vent corridors (Anderson and others, 1999, p. 27). Secondary mineralization is uncommon in these rocks, except where carbonates fill vesicles near the land surface (Nace and others, 1975, p. 13).
Results of single-well aquifer tests in 67 wells with perforated or open intervals only in hydrogeologic unit 1 indicated that the hydraulic conductivity of these rocks ranges from about 0.01 to 24,000 ft/d (table 2), more than six orders of magnitude. Almost two-thirds of these estimates were larger than 100 ft/d and about one-third were larger than 1,000 ft/d. Estimates larger than 100 ft/d primarily are associated with interflow zones of thin pahoehoe flows (Anderson and others, 1999, p. 27). Many estimates of hydraulic conductivity smaller than 100 ft/d may be associated with dikes (Anderson and others, 1999, p. 27).
Three previous studies described specific yield, porosity, and effective porosity applicable to hydrogeologic unit 1. In the RASA study, Garabedian (1992, p. 44-46) used an average specific yield of 0.05 for all types of basalt and silicic volcanic rocks and 0.20 for all types of sediment for the uppermost 200 ft of the eastern SRP aquifer. Ackerman (1995, p. 10) used a range of effective porosity from 0.10 to 0.25 for simulation of advective transport in the uppermost 200 ft of the regional aquifer system. The calibration value for this simulation, 0.21, was based on the travel time of iodine-129 (129I) from disposal well CPP 3 at the INTEC to well USGS 11 near Big Southern Butte (fig. 1) (Ackerman, 1995, p. 22). Knutson and others (1992, p. 4-21 and fig. 4-10) determined the porosity of 1,504 core samples from the unsaturated zone of hydrogeologic unit 1 near the RWMC. The porosities ranged from 0.01 to 0.43 with 90 percent of values between 0.05 and 0.27.
Hydrogeologic unit 2 consists of 7 basalt-flow groups, 10 sedimentary interbeds, and includes basalt-flow group I, a major basalt-flow group within composite unit 7 (fig. 11) (Anderson and Liszewski, 1997, p. 16, figs. 9-15). These massive, less densely fractured basalt flows are mostly 21- to 37-ft thick (table 3) tube-fed pahoehoe flows having thick, massive interiors (Anderson and others, 1999). Mean thickness of the basalt flows in hydrogeologic unit 2 is 29 ft (table 3), or nearly one-and-one-half times the average thickness of the basalt flows composing hydrogeologic unit 1. As a result of the larger average thickness of basalt flows, the basalt of hydrogeologic unit 2 is probably more massive, less densely fractured, and has fewer interflow zones than hydrogeologic unit 1.
The massive basalt of hydrogeologic unit 2 includes basalt-flow group I, one of the thickest and most extensive flow groups near the INL (Anderson, 1991, p. 22; Wetmore and others, 1997, p. 50, table 1). The interpreted thicknesses of the other basalt-flow groups in hydrogeologic unit 2 were based entirely on natural gamma logs and are less certain than the interpretations for basalt-flow group I. Basalt-flow group I and the overlying HI interbed (a layer of clay and silt) underlie all but the northern and extreme southeastern parts of the INL (Anderson, 1991, p. 21-22; Anderson and others, 1997, p. 19). The thickness of composite stratigraphic unit 7 ranges from 0 to 409 ft, averages 266 ft in the 15 wells that fully penetrate to its base, and is greatest in wells at the TRA near the exposed vents of basalt-flow group I. Correlations of composite stratigraphic unit 7 were primarily based on natural-gamma logs from uncored boreholes because only five coreholes within a 3-mi radius of the INTEC are known to penetrate this stratigraphic unit.
Hydraulic conductivity and porosity of hydrogeologic unit 2 generally are less than hydrogeologic unit 1 because of the thick, massive interiors and fewer interconnected interflow zones in the massive, less densely fractured basalt. Other factors that probably modify the hydraulic conductivity and porosity of the fractured basalt include sediment-filled interflow zones and near-vent deposits, dikes, fissures, and tension cracks in volcanic rift zones and vent corridors (Anderson and others, 1999, p. 27). Secondary mineralization is uncommon in these rocks, except where carbonates fill vesicles near the land surface (Nace and others, 1975, p. 13).
Estimated hydraulic conductivity of hydrogeologic unit 2 ranged from 6.5 to 1,400 ft/d for single-well aquifer tests in four wells with perforated intervals only in hydrogeologic unit 2 (table 2). This range of conductivities is similar to conductivity ranges estimated for massive basalt of the Columbia Plateau (Anderson and others, 1999, p. 22 and 27), results of flowmeter tests in two wells penetrating part of unit 2 (Morin and others, 1993, p. 23 and 26), and numerical model simulations of aquifer tests that include unit 2 (Frederick and Johnson, 1996, p. 63, layer 3).
Porosity of the massive basalt of hydrogeologic unit 2 probably is within the lower end of the range estimated for that of the densely fractured basalt of hydrogeologic unit 1, 0.05 to 0.27. Knutson and others (1992, p. 4-21) reported a median porosity of 0.11 for hundreds of nonvesicular basalt cores, a measure that may approximate the porosity of massive basalt because it does not include the effects of porous interflow zones.
Intermediate-age rocks of hydrogeologic unit 3 underlie younger rocks in the model area, except for the area northwest of the Big Lost River (Anderson and Liszewski, 1997, fig. 30) (fig. 1) where the intermediate-age rocks are not present. The intermediate-age rocks constitute the full thickness of the aquifer in the northern part of the INL and in the southwestern part of the model area (figs. 9 and 10). Hydrogeologic unit 3 consists of 41 basalt-flow groups, 1 andesite-flow group, and 26 sedimentary interbeds (fig. 11) (Anderson and Liszewski, 1997, p. 17-18, figs. 9-15). Sediment content is greatest in a sedimentary trough, known as the Big Lost Trough, in the northern part of the model area. Average thickness of fractured basalt flows in hydrogeologic unit 3 is 23 ft and 50 percent of the flows are between 15- and 28-ft thick (table 3).
Basalt flows of hydrogeologic unit 3 are slightly-to-moderately altered (Whitehead, 1992, p. 10, table 3; Lanphere and others, 1994, p. 22-39; Fromm and others, 1994; Anderson and Liszewski, 1997, p. 28). Detailed analysis of petrography of cores from wells near TAN indicates that intermediate-age rocks near TAN contain secondary pore-filling minerals such as calcite (Lanphere and others, 1994, p. 22-39). Calcite also is commonly present in the vesicles and fractures of these rocks. This alteration and the larger average thickness of the basalt flows, compared with the thickness of basalt flows in hydrogeologic unit 1 (Whitehead, 1992, p. 11-13), result in fewer interflow zones and a smaller hydraulic conductivity and porosity for hydrogeologic unit 3 than for hydrogeologic unit 1.
Results of single-well aquifer tests in 14 wells with perforated or open intervals only in hydrogeologic unit 3 indicated that the hydraulic conductivity of these rocks ranges from about 0.32 to 24,000 ft/d (table 2). However, a comparison of 24 hydraulic conductivity estimates near the INTEC with 68 estimates near TAN indicate that the average hydraulic conductivity of intermediate-age rocks near TAN is about one order of magnitude smaller than younger rocks (hydrogeologic units 1 and 2 undifferentiated) near the INTEC (John Welhan, Idaho State Geological Survey, written commun., 1999). Welhan calculated a median and geometric mean hydraulic conductivity of 500 and 130 ft/d for younger rocks near the INTEC and 30 and 20 ft/d for the intermediate-age rocks near TAN. In the analysis of the regional aquifer system, Garabedian (1992, p. 42) accounted for this difference by decreasing the hydraulic conductivity of lower model layers by as much as two-thirds.
Porosity of hydrogeologic unit 3 probably is within the lower end of the range of that estimated for the densely fractured basalt of hydrogeologic unit 1, 0.05 to 0.27. Median values of porosity reported for 10 nonvesicular and 10 vesicular basalt cores from the intermediate-age rocks at TAN were about 0.05 and 0.08, respectively (Allan Wylie, Idaho Water Resources Research Institute, written commun., 2000). Median values of porosity reported for hundreds of nonvesicular and vesicular basalt cores from the younger rocks (hydrogeologic units 1 and 2) at and near the RWMC were 0.11 and 0.22, respectively (Knutson and others, 1992, p. 4‑21). Although these values do not include porosity of interflow zones, they indicate that porosity of intermediate-age rocks may be smaller than that of younger rocks.
Older rocks (Tb, Tsv, and QTb in fig. 4) underlie intermediate-age rocks throughout most of the model area and contain intensely altered basalt (which generally is equivalent to the Glenns Ferry Formation), Miocene rhyolitic ash-flow tuffs (figs. 4 and 11), and interbedded sediment. Only a few estimates of hydraulic conductivity and porosity were available for the older rocks. Based on analyses from four aquifer tests, Mann (1986, p. 21) reported a range of 0.002 to 0.03 ft/d for hydraulic conductivity of altered basalt, interbedded sediment, and rhyolitic ash-flow tuffs that make up the older rocks. Nace and others (1959, table 9) reported two measurements of porosity for older basalt, both about 0.19, at a spring in the discharge area of the eastern SRP aquifer. This value is greater than the median porosity, about 0.09, of 17 cores of altered basalt obtained from the older rocks at well WO2 (fig. 1), about 3 mi east of the INTEC (Carroll Knutson, EG&G Idaho, Inc., written commun, 1992).
Rhyolite domes in the model area are clustered in two areas (fig. 9; rhyolite domes A, B, C, and an unnamed dome between B and C that is not shown in fig. 9). Rhyolite domes are vertical plug-like masses (Qsv in fig. 4) interpreted to penetrate a large thickness of the younger rocks and intermediate-age rocks (Kuntz and Dalrymple, 1979, p. 30-34; Spear and King, 1982, p. 396-400; Kuntz and others, 1994; Hughes and others, 1999, fig. 16; McCurry and others, 1999, p. 170-174). Big Southern Butte is the largest of four rhyolite domes. Other domes, Middle Butte, East Butte, and an older unnamed dome of small surficial extent between Middle Butte and East Butte, are near the southeast boundary of the model area.
Hydraulic properties of rhyolite domes have not been measured; however, based on rock characteristics, water-table contours (Bartholomay and others, 1997, fig. 9), and well Corehole 1 (CH1 in fig. 1) that penetrates the saturated part of the unnamed dome (Kuntz and Dalrymple, 1979, fig. 5; Morse and McCurry, 1997, fig. 2; McCurry and others, 1999, fig. 3), rhyolite domes probably have low permeability and may have hydraulic properties similar to those of the massive basalt of hydrogeologic unit 2. A temperature log of Corehole 1 (Morse and McCurry, 2002, fig. 2) indicates that most of the unnamed dome is in hydraulic contact with the cold water of the aquifer. However, inferred water-table contour deflections around this dome and Middle and East Buttes (fig. 12) probably indicates that the hydraulic conductivity of these rocks is smaller than that of adjacent fractured basalt. The lack of a similar deflection of contours near Big Southern Butte probably indicates that this dome is fractured at depth or may lie above the water table. If the hydraulic conductivity and porosity of rhyolite domes are similar to that of the massive basalt of hydrogeologic unit 2, then the conductivity probably ranges from about 6.5 to 1,400 ft/d and the porosity is probably within the lower end of the 0.05 to 0.27 range of porosity estimated for hydrogeologic unit 1. In the RASA study, Garabedian (1992, p. 44-46, tables 19 and 20) used a model-calibrated value of 0.65 ft/d for hydraulic conductivity and a value of 0.05 for specific yield of silicic volcanic rocks.
Sediment in the model area is interbedded with basalt and deposited in four overlapping depositional environments in younger rocks, intermediate-age rocks, and older rocks (Whitehead, 1992; Kuntz and others, 1994; Anderson and others, 1996a, b). Sediment deposited in the first depositional environment consists of thin layers of windblown sediment that are present throughout most of the model area (Nace and others, 1975, p. 35; Anderson and others, 1996b, p. 3). Sediments from the second depositional environment range from fluvial, sandy gravel in stream channels to finer-grained clayey silt in terminal playas at the distal ends of river systems (Nace and others, 1975, p. 19-27; Gianniny and others, 1997, p. 31). Mainly thick layers of clay and silty clay lacustrine deposits in an area of closed topographic depressions at and near Mud Lake (Stearns and others, 1939, p.17, 39; Spinazola, 1994, p. 10) constitute sediment from the third depositional environment. Alluvial deposits along and near the mouths of tributary valleys and adjacent mountain fronts mainly constitute sediment from the fourth depositional environment.
Sedimentary deposits along and near the channel, floodplain, sinks, and playas of the Big Lost River are referred to as the Big Lost Trough (Gianniny and others, 1997, p. 31). The definition of the Big Lost Trough is expanded in this report beyond that of Gianniny and others (1997, p. 31) to include the coalescing sedimentary deposits along and near the channel, floodplain, sinks, and playas of the Big Lost River, Little Lost River, Birch Creek, Camas Creek, and Mud Lake (fig. 12). Two interpretations of the Big Lost Trough are shown in figure 12 to account for differences in methods, data, and interpolation techniques used to evaluate the distribution of sediment. Figure 12A shows an area, interpreted from drill logs (Whitehead, 1992, pl. 5), where sediment was estimated to compose from 100 to 999 ft of the stratigraphic section (including the older rocks). Figure 12B shows an area where sediment was estimated to compose more than 11 percent of the stratigraphic section (excluding the older rocks) based on interpretations from the geologic map of the INL (Kuntz and others, 1994) and natural-gamma logs from wells drilled to depths of more than 300 ft below land surface (Anderson and others, 1996a, b). Sedimentary deposits compose from less than 1 percent to 50 percent of the rocks penetrated by wells in the model area. Although sediment is present throughout the stratigraphic section of the model area, it is not represented with separate hydrogeologic units, but is included in younger rocks (hydrogeologic units 1 and 2), intermediate-age rocks (hydrogeologic unit 3), and older rocks.
Estimates of hydraulic conductivity and porosity measured by or estimated for previous studies reflect the wide range of characteristics of the sedimentary interbeds beneath the INL. For example, laboratory determinations of saturated hydraulic conductivity for 109 core samples from selected sedimentary interbeds at and near the RWMC range from about 0.000032 to 240 ft/d (McElroy and Hubbell, 1990; Winfield, 2005). Reported textures of these samples range from clay to sand. By comparison, for the RASA study, Garabedian (1992, tables 19 and 20) used model-calibrated ranges of hydraulic conductivity from 0.0033 to 0.65 ft/d for silt and clay, 0.65 to 5,200 ft/d for sand, and 33 to 17,000 ft/d for sand and gravel. Robertson (1977, p. 21) reported a range in porosity of from 0.25 to 0.45 for selected surficial sediments. Laboratory determinations of porosity for 109 core samples from selected sedimentary interbeds at and near the RWMC and INTEC range from about 0.33 to 0.73 (McElroy and Hubbell, 1990, table 8; Perkins and Nimmo, 2000, table 1; Perkins, 2003, table 1; Winfield, 2003, table 1; Winfield, 2005, table 2). Estimates of effective porosity used for modeling contaminant movement through the unsaturated zone at the INL were 0.30 and 0.35 (Robertson, 1977, p. 31, 45). In the RASA study, Garabedian (1992, p. 44-46) specified a value of 0.20 for the specific yield of sediments.
Hydraulic properties of interbedded sediment probably significantly influence the hydraulic properties of basalt in the eastern SRP aquifer. However, because hydraulic conductivity and porosity for sediment were available for only a few sedimentary interbeds in the model area and were not available for complete interflow zones, the influence of sediment on hydraulic properties of hydrogeologic units was not directly included in the conceptual model. For instance, in areas where large amounts of sediment were deposited, such as in the Big Lost Trough, hydraulic conductivity of basalt interflow zones probably is greatly reduced because sediment fills cracks, joints, fissures, and fractures, reducing the original porosity of the basalt and impeding ground-water flow. The effects of sedimentary deposits on hydraulic conductivities of hydrogeologic units could be estimated with a numerical model.
The model area, in the west-central part of the eastern SRP aquifer, is bounded by three physical and three artificial hydrologic boundaries. Physical boundaries are the water-table boundary, base of the aquifer, and northwest mountain-front boundary. Artificial boundaries are the northeast boundary, southeast-flowline boundary, and southwest boundary (fig. 13). Artificial boundaries were used for the model because using the boundaries of the eastern SRP aquifer would produce a model area much larger than that required for the purposes of this study and would include more area with little available information. Artificial boundaries were located (1) to include within the model area the part of the eastern SRP aquifer containing detectable concentrations of contaminants derived from wastewater discharged at the INL, and (2) in areas where information was available to define and characterize the boundary and flows across the boundary.
The water-table boundary is represented by the contact between the unsaturated zone and the saturated zone. The thickness of the unsaturated zone ranges from about 200 to 1,000 ft (fig. 10). The altitude of the water table decreases about 500 ft across the length of the model area, from 4,600 ft at the northeast boundary of the model to about 4,100 ft at the southwest boundary of the model (fig. 9). The water-table boundary rises or declines in response to climatic- or irrigation-induced changes in recharge to the aquifer.
The base of the aquifer is interpreted as the contact between the older rocks and the overlying intermediate-age rocks. This contact generally is characterized by an abrupt downward increase in intensely altered basalt and interbedded sediment (Mann, 1986, p. 4-5; Whitehead, 1992, p. 10; Morse and McCurry, 1997 p. 6-7; Anderson and Liszewski, 1997, p. 28). In seven wells and coreholes, this contact also coincides with a change from convective-dominated to conductive-dominated heat flow and a decrease in hydraulic conductivity and porosity (Mann, 1986, p. 21; Morse and McCurry, 1997, p. 6-7; Welhan and Wylie, 1997, p. 99; Morse and McCurry, 2002, p. 222; Mazurek and others, 2004). Collectively, these factors indicate that water circulation is limited in older rocks.
Direct evidence indicating the location of the base of the aquifer was limited to data from 13 deep wells and coreholes (Mann, 1986, p. 21; Whitehead, 1992, fig. 8 and pl. 6; Morse and McCurry, 1997, p. 4-7, fig. 2; Anderson and Liszewski, 1997, table 3; Mazurek and others, 2004). Morse and McCurry (1997, 2002) correlated the base of the aquifer with an increase in alteration of basalts and temperature gradient in six deep wells where temperature was measured (Blackwell and others, 1992). Mazurek and others (2004) made the same correlation based on data from a more recently drilled deep borehole. Morse and McCurry (1997) estimated depth to the base of the aquifer in the central part of the INL to be about 1,340 ft. Anderson and Liszewski (1997, p. 4), on the basis of stratigraphic data from 10 deep wells and coreholes in the western half of the INL, suggested that the base most likely coincides with the top of a thick and widespread layer of clay, silt, sand, and altered basalt that is consistent in age with rocks of the Glenns Ferry Formation. They reported a depth to the base of the aquifer in the western half of the INL ranging from 815 to 1,710 ft below land surface and a saturated thickness ranging from 445 to 1,200 ft.
Indirect evidence indicating the location of the base of the aquifer included regional surface-based electrical-resistivity surveys (Whitehead, 1992, pl. 6) and regional gravity surveys (Whitehead, 1992, fig. 8). Based on surface-based electrical-resistivity soundings and drillhole data, Whitehead (1986, sheet 2, section H-H´) estimated that (1) depth to the base of the aquifer could exceed 3,000 ft in the eastern half of the INL, and (2) saturated thickness of the aquifer could exceed 2,500 ft in the eastern half of the INL and 4,000 ft in the southwestern part of the model area. Whitehead (1986, sheet 2) reported saturated thickness using contour intervals of 500 ft; therefore, the uncertainty of these estimates is at least ±250 ft. However, because electrical-resistivity surveys tend to overestimate basalt thickness (Zohdy, 1974, p. 32), these estimates of depth to the base of the aquifer and saturated thickness may be high.
Whitehead’s (1986, sheet 2) data were used to calculate the depth to the base of the aquifer, rather than data from the 13 deep wells and coreholes, because Whitehead’s data provide coverage over the entire model area. Using Whitehead’s (1986, sheet 2) data for thickness of Quaternary basalt and sediment and a digital elevation model, depth to the base of the aquifer was calculated using a geographic information system. However, depth to the base of the aquifer in the area mountainward of the Big Lost River was subsequently adjusted to a lesser depth based on stratigraphic interpretations that indicate the aquifer is thinner in this area (Anderson and Liszewski, 1997, figs. 29, 30). Interpreted depth to the base of the aquifer in the model area ranges from about 700 to 4,800 ft. Depth is least near the northwest mountain-front boundary, increases to the east, and is greatest along the southeast-flowline boundary (fig. 14). A shaded relief map (fig. 15) of the contact between the intermediate-age rocks of hydrogeologic unit 3 and the older rocks that form the base of the aquifer indicates that the inferred three-dimensional topography along the base of the aquifer is irregular.
The northwest mountain-front boundary is defined by the edge of the mountain fronts of the Pioneer Mountains, Lost River Range, Lemhi Range, and Bitterroot Range and the mouths of the Big Lost River, Little Lost River, and Birch Creek valleys (fig. 1). This boundary coincides with part of the northwest boundary of the eastern SRP as described in the RASA study (Garabedian, 1992). Aquifer thickness along the northwest mountain-front boundary ranges from about 200 to 500 ft.
The northeast boundary is upgradient of the INL in an area where the hydraulic gradient increases sharply (figs. 3 and 9). The hydraulic gradient in this area ranges from 27 to 60 ft/mi and averages about 36 ft/mi (fig. 12A) (Lindholm and others, 1988). These steep gradients coincide with changes in aquifer transmissivity near Mud Lake where basalt interfingers with less transmissive layers of sediment (Crosthwaite, 1973, p. 6; Lindholm, 1996, p. 22) and probably are caused by a decrease in hydraulic conductivity associated with these fine-grained sediments (Nace and others, 1959, p. 151-2; Mundorff and others, 1964, p. 133; Lindholm and others, 1988). This low-transmissivity zone impedes horizontal flow (Crosthwaite, 1973, p. 11; Spinazola, 1994, p. 29) and results in a large downward vertical component of flow. Aquifer thickness along the northeast boundary ranges from about 500 to 800 ft.
The southeast-flowline boundary is near the central axis of the eastern SRP (figs. 3 and 13), and the aquifer thickness along this boundary ranges from about 800 to 4,000 ft. The boundary is represented as a southwest trending flowline derived from pathline analyses (Ackerman, 1995) of the steady-state, three-layer RASA model (Garabedian, 1992). Ackerman’s model computed three-dimensional pathlines that originated as particles introduced near the water table. Particles introduced at the water table upgradient of the model area followed nearly parallel pathlines along the model area’s southeastern boundary (Ackerman, 1995, fig. 8). These simulations showed that, although most flow is horizontal and contained in the uppermost 500 ft of the aquifer, particles that penetrate deeper into the aquifer also follow nearly parallel pathlines along this boundary. The sustained parallelism of these pathlines implies that little or no lateral divergence of flow occurs at depth along the central axis of the eastern SRP aquifer (fig. 3) east of the INL. Projections of pathlines to the water table indicated that particle pathlines at depth are nearly parallel to flowlines oriented perpendicular to water-table contours in the central part and along the central axis of the eastern SRP aquifer east of the INL (Robertson and others, 1974; Barraclough and Jensen, 1976; Barraclough and others, 1981; Lewis and Jensen, 1985; Pittman and others, 1988; Orr and Cecil, 1991; Ackerman, 1995, fig. 8; and Bartholomay and others, 1995). Therefore, the pathline representing the southeast boundary of the model area can be approximated as a flowline in a vertical subsurface projection.
The southwest boundary is located so the model includes areas of the aquifer downgradient of where contaminants were detected and attributed to wastewater discharged at the INL (fig. 2). No physical hydrologic features are present that could be used as a model boundary in this area. Consequently, an artificial boundary was located along a gridline from the RASA model (fig. 16) about 25 mi downgradient of the southwestern extent of chlorine-36 (36Cl) detections in the aquifer (Beasley, 1995, appendix 2). Few wells penetrate the aquifer in this area, so the subsurface geology and hydraulic properties along the southwest boundary are not well known. Aquifer thickness along this boundary ranges from about 500 to 3,000 ft.
Inflows, outflows, and fluxes (flow per unit area) across the model boundaries (fig. 17 and table 4) were estimated on the basis of hydrologic data collected and interpretive studies done during the past 50 years and application of Darcy’s Law to physical and hydrologic features along these boundaries. The past 50 years represent a time when the hydrologic conditions of the system were relatively stable compared with conditions in the first 50 years of the twentieth century, when changes in irrigation practices, climatic variations, and surface-water diversions placed greater stresses on the system (Garabedian, 1992, p. 31, 50). Inflows and outflows across model boundaries were conceptualized as spatially uniform or nonuniform, and as temporally constant or variable (table 4). These boundary characteristics represent varying degrees of simplification that reflect (1) physical features of the boundary (for example, the tributary-valleys along the northwest mountain-front boundary), (2) hydraulic features of the boundary (for example, the water-table gradients along the northeast boundary), (3) data limitations that do not justify characterization of flow across the boundary in greater detail (for example, the base of the aquifer), (4) proximity of flow across the boundary to areas of known aquifer contamination (for example, industrial wastewater return flows), and (5) the magnitude and relative importance of the boundary flux to contaminant transport (for example, precipitation and streamflow-infiltration recharge).
Streamflow infiltration and underflow from the tributary streams and valleys (Big Lost River, Little Lost River, and Birch Creek), respectively, provide recharge to the eastern SRP aquifer. In the conceptual model, infiltration recharge from the tributary streams was treated as flow across the water-table boundary, and underflow from the tributary valleys was treated as flow across the northwest mountain-front boundary.
Inflows and outflows across the water-table boundary (fig. 13) were treated implicitly1 in the conceptual model and comprise five components: (1) diffuse areal precipitation recharge; (2) infiltration recharge from streams originating in tributary valleys; (3) ground-water withdrawals for irrigation and industrial use; (4) wastewater return flows; (5) and irrigation-infiltration recharge. Release of water from storage, which is an internal flow, is discussed with flows across the water-table boundary. Individual flow components may be small or zero during some periods, but may be important because of their proximity to wastewater discharged in the eastern SRP aquifer at the INL.
1 Flow through the unsaturated zone is treated as time-averaged (annual and seasonal) net-infiltration recharge to the aquifer. Downward flow in response to changes in the saturation-matric potential and saturation-hydraulic conductivity relations embodied in the Richards’ equation for unsaturated flow is assumed to be included in inflow components of the water budget that are areally uniform (precipitation) or areally nonuniform (streamflow infiltration, industrial wastewater return flow, and irrigation infiltration) and temporally constant (precipitation) or temporally variable (streamflow infiltration, industrial wastewater return flow, and irrigation return flow). Typical unsaturated-zone flow processes, such as moisture depletion from evapotranspiration, downward percolation and lateral redistribution, wetting and drying, and perching, are not treated as distinct flow processes.
Although the quantity, temporal variability, and spatial distribution of precipitation recharge in the model area are not well defined, studies indicate that the net effect of this recharge probably is small and is of less importance than that of other recharge sources. The effect of precipitation recharge in the model area has not been discernable in water-level measurements. Mundorff and others (1964, p. 184) estimated precipitation recharge for two areas of different soil cover in the model area. Precipitation recharge was 0.02 ft/yr in the central part of the eastern SRP. Precipitation recharge was 0.3 ft/yr near Craters of the Moon National Monument (fig. 1), where precipitation is greater and fractured basalt and rubble cover the surface. Kjelstrom (1995, p. 11) reported an average precipitation recharge of 0.08 ft/yr based on a basin-wide water budget for the eastern SRP. Based on 36Cl and tritium (3H) profiles and neutron logging, Cecil and others (1992, p. 713) estimated that net precipitation recharge at the INL ranges from 2 to 5 percent of mean annual precipitation. Precipitation at the INL is about 0.7 ft/yr (Clawson and others, 1989, tables D-1 and D-2; Goodell, 1988, fig. 5), and 2 to 5 percent of this value would be from 0.01 to 0.04 ft/yr. Using a maximum recharge rate of 0.04 ft/yr for the 1,940 mi2 of the model area, maximum mean annual precipitation recharge across the water table is about 70 ft3/s. A uniform distribution of 70 ft3/s of precipitation recharge over the model area would result in an infiltration flux of about 0.04 (ft3/s)/mi2. This very low areal flux indicates that the effect of precipitation recharge on ground-water flow directions and velocities is probably very small. In areas where precipitation accumulates as runoff into small, closed basins, numerous sedimentary interbeds in the thick, unsaturated zone (200-1,000 ft) serve to intercept and laterally redistribute downward percolation. This redistribution tends to offset effects of local runoff accumulation and maintains an approximately uniform and constant rate of recharge to the aquifer.
Streamflow onto the INL is intermittent and nearly all streamflow infiltrates through the thick, unsaturated zone and recharges the eastern SRP aquifer. Streams tributary to the eastern SRP near the INL originate in mountain ranges north and west of the model area (fig. 15) and include the Big Lost River, the Little Lost River, and Birch Creek. Routine streamflow measurements were made by the USGS at several streamflow-gaging stations on these streams for more than 50 years.
Because of its large flux and proximity to known sources and areas of contamination in the aquifer, streamflow-infiltration recharge from the Big Lost River is considered more significant to contaminant transport than the other components of inflow across the water-table boundary. During years with sufficient precipitation, the Big Lost River flows onto the eastern SRP near Arco and terminates on the INL in a series of four playas southwest of TAN (Kjelstrom and Berenbrock, 1996). Water from the river is stored in Mackay Reservoir (capacity 44,370 acre-ft; Bennett, 1990) (fig. 1), about 40 mi upstream of the INL. In most years, much of the water in the river downstream of the reservoir is diverted for irrigation or infiltrates through the riverbed before reaching Arco (Kjelstrom and Berenbrock, 1996). The average annual streamflow of the Big Lost River near Arco for the period of record available between 1946 and 2003 was 95 ft3/s (fig. 18, gaging station 13132500); (Brennan and others, 2004, p. 196). However, streamflow fluctuates greatly in response to short-term (3 to 8 years) climate cycles. During dry climate cycles, streamflow in the Big Lost River near Arco frequently is zero (fig. 18). In 1984, however, during a wet climate cycle, the mean annual streamflow at the gaging station near Arco was 488 ft3/s (fig. 18), or more than 20 percent of the estimated 1980 outflow across the southwest boundary of the model area (table 4).
Flow from the Big Lost River is diverted to spreading areas southwest of the RWMC near the southwest boundary of the INL (fig. 1) to prevent flooding of site facilities. The spreading areas were constructed in 1958, first used in 1965, and enlarged in 1984 (Bennett, 1990, p. 6). Between 1965 and 2002, about 47 percent (48 ft3/s, table 5) of the streamflow at the INL diversion [(103 ft3/s, table 5), estimated from the combined streamflow for gaging stations INL diversion at head near Arco (gaging station 13132513, fig. 1) and Big Lost River below diversion near Arco (gaging station 13132520, fig. 1)] was diverted to the spreading areas. Because the spreading areas were used for flood control, during wet climate cycles an even larger percentage of streamflow was diverted to the spreading areas. For instance, in 1984, the year of highest recorded flows, about 83 percent (379 ft3/s, fig. 18, and table 5) of streamflow at the INL diversion was diverted to the spreading areas. The use of the spreading areas has changed the location of much of the streamflow-infiltration recharge to the aquifer. Before 1965, the Big Lost River recharged mostly in the sinks and playas in the west-central part of the INL; currently, when flow extends onto the plain, much of it recharges in the spreading areas near the southwest boundary of the INL.
Streamflow onto the INL is spatially concentrated in the channel, sinks, playas, and spreading areas (about 5 mi2 total) (fig. 1). Infiltration losses from the river channel were estimated to be from about 1 to 4 (ft3/s)/mi between Arco and the Big Lost River Sinks to as much as 28 (ft3/s)/mi from the Big Lost River Sinks (Bennett, 1990, p. 24 and 26). Infiltration and evaporation rates from the Big Lost River playas were estimated to be 0.16 to 0.37 ft/d and 0.01 to 0.02 ft/d, respectively, and infiltration rates from Big Lost River spreading areas A and B were estimated to be 0.7 and 2.6 ft/d, respectively (Barraclough and others, 1967, p. 24). Owing to the much larger rates of infiltration than evaporation, all streamflow in the model area was assumed to infiltrate to the aquifer.
Large amounts of streamflow during wet climate cycles provide a source of episodic streamflow-infiltration recharge to the aquifer. Average infiltration flux from streamflow is about 20 (ft3/s)/mi2, which is more than two orders of magnitude larger than precipitation recharge per unit area. However, during 1984, a year of record-high diversions to the spreading areas, average infiltration flux at the spreading areas was 170 (ft3/s)/mi2 and monthly infiltration fluxes ranged from about 80 to about 360 (ft3/s)/mi2 [calculated using mean monthly streamflow to the spreading areas (Bennett, 1990, table 4) and a combined area for spreading areas A, B, C, and D of 2.3 mi2]. Maximum monthly infiltration flux was nearly four orders of magnitude larger than precipitation recharge per unit area.
Data were not adequate to estimate infiltration recharge from the Little Lost River to the model area; however, most streamflow either infiltrates or is diverted for irrigation upstream of the INL (Kjelstrom and Berenbrock, 1996, p. 3). When the Little Lost River flows onto the eastern SRP, it terminates in an area commonly referred to as the Little Lost River Sinks, near Howe, just east of the northwest mountain-front boundary of the model area (fig. 1). Stone and others (1992, p. 20-21) reported a mean annual streamflow of 77 ft3/s for the combined periods of record 1941–81 and 1986–90 at a streamflow-gaging station on the Little Lost River about 7 mi northwest of Howe (streamflow-gaging station 13119000, fig. 1).
The amount of infiltration from Birch Creek streamflow to the eastern SRP aquifer is not well documented because streamflow-gaging stations upstream of the INL boundary were discontinued or operated intermittently (Kjelstrom and Berenbrock, 1996, p. 7-8). Flow from Birch Creek is diverted upgradient of the model area to a hydroelectric power-generating facility several miles east of Birch Creek. Return flows as large as 46 ft3/s from the power facility (Ted S. Sorenson, Sorenson Engineering, written commun., March 2001) are diverted to the northeast for irrigation use for 6 months each year. During the nonirrigation season, this water flows onto the northeastern part of the INL, where it may infiltrate to the aquifer through diversion channels and gravel pits northwest of TAN (Kjelstrom and Berenbrock, 1996, p. 5). No flow has reached the Birch Creek Sinks near TAN since 1969, when accumulation of water in the playas prompted construction of channels to divert water away from TAN (Kjelstrom and Berenbrock, 1996, p. 5).
Total ground-water withdrawals and industrial wastewater return flows at the INL in 1980 were 45 and 6 ft3/s, respectively (fig. 17, table 4). Irrigation-infiltration recharge in the model area was estimated at about 24 ft3/s. Ground-water withdrawals in the model area include those for industrial use at onsite facilities and those for irrigation use offsite. Industrial wastewater return flows include wastewater discharged in (1) disposal ponds or ditches, where it infiltrates into the unsaturated zone and the aquifer, and (2) injection wells completed in the aquifer.
Industrial ground-water withdrawals and wastewater return flows ranged from zero before the INL was established to a maximum in 1974 of about 12 ft3/s for withdrawals and 7.3 ft3/s for return flows (White, 1975). The volume of industrial ground-water withdrawals from the aquifer and wastewater discharged in infiltration ponds and wells varied over the years in response to the amount of industrial activity at INL facilities and changes in wastewater-disposal methods (Appendix A). In 1980, total industrial ground-water withdrawals at the INL were about 8.2 ft3/s, the volume of wastewater discharged to ponds and ditches was about 2.5 ft3/s, and the volume of wastewater discharged to injection wells was about 3.4 ft3/s (Batchelder, 1981). Local effects of these withdrawals and returns at INL facilities were not expected to have long-term effects on aquifer response beneath the central part of the INL; however, withdrawals and returns at TRA and INTEC could have significant short-term effects on contaminant migration. Most contaminants in the aquifer beneath the INL originate at these two facilities and more than 50 percent of industrial ground-water withdrawals and 50 percent of industrial wastewater returns were pumped from or injected into the eastern SRP aquifer at these two facilities.
Ground-water withdrawals for irrigation increased substantially during the 1940s and continued to increase until the 1970s (Kjelstrom, 1995, p. 5). For 1980, estimated irrigation withdrawals in the model area are 37 ft3/s. Ground water used for irrigation in the model area is withdrawn from the aquifer primarily near Mud Lake, but smaller quantities also are withdrawn near Howe (fig. 1) (Garabedian, 1992, pl. 3). Spinazola (1994, p. 33-38) calculated withdrawals near Mud Lake based on electrical-power-consumption records, irrigation-system characteristics, and hydrologic measurements made between 1983 and 1990, and determined that ground-water withdrawals for irrigation in the Mud Lake area in 1983 (similar to those in 1980) were about 31 ft3/s. Based on estimates of irrigated acreage (DeWayne McAndrew, U.S. Bureau of Reclamation, written commun., 1996) and the volume of water used per acre in the Mud Lake area, an annual average of about 6 ft3/s is withdrawn for irrigation near Howe.
Estimates of irrigation-infiltration fluxes for the SRP were reported by Goodell (1988) and Spinazola (1994). Goodell (1988, p. 48, table 15) estimated average infiltration fluxes for ground- and surface-water irrigation for the whole SRP to be 0.62 and 4.1 (ft3/s)/mi2, respectively. The infiltration flux for surface-water irrigation estimated by Goodell (1988, p. 48, table 15) includes conveyance losses between points of withdrawal and application that are outside the model area. Spinazola (1994, p. 29-33, fig. 38) estimated infiltration fluxes at and north of the model area to be 0.15 (ft3/s)/mi2 from sprinkler-applied irrigation (primarily from ground-water sources) and 0.71 (ft3/s)/mi2 from ditch-applied irrigation (from surface-water and diverted ground-water sources). Spinazola’s (1994) estimates did not have to be adjusted for conveyance losses, which were estimated separately from irrigation-infiltration fluxes, and included estimates of consumptive use for the location where irrigation was applied. Using a geographic-information-system coverage of irrigated acreage derived from 1987 aerial photographs and field checked in 1992 (DeWayne McAndrew, U.S. Bureau of Reclamation, written commun., 1996), about 40 and 26 mi2 of the model area was sprinkler- and ditch-irrigated land, respectively. Based on these estimates of irrigated areas and Spinazola’s (1994) estimates of irrigation-infiltration fluxes, about 24 ft3/s of irrigation-infiltration recharged to the model area.
Release of water from storage in the model area for 1980 was estimated using measured changes in water levels from October 1979 to October 1980 (Garabedian, 1986, pl. 1D). Using RASA derived storage coefficients ranging from 0.05 to 0.2, the calculated net gain to the aquifer from a decrease in storage during this period was 60 to 240 ft3/s. Using storage coefficients of 0.1 for areas dominated by basalt and 0.2 for areas dominated by sediment, the net gain to the aquifer from a decrease in storage during this period was 80 ft3/s.
Because of its low flux and distance from known areas of contamination in the aquifer, the base of the aquifer, for purposes of contaminant transport modeling, can be treated as a no-flow boundary. Upward flow across the base of the aquifer was estimated to be about 44 ft3/s based on data from a few wells that penetrate the base of the aquifer. Temperature data from six deep boreholes indicate that the thermal gradient across the base of the aquifer changes from convective-dominated heat flow in the aquifer to conductive-dominated heat flow in the rocks below the aquifer (Morse and McCurry, 2002, p. 215). This change in thermal gradient implies that flow across this boundary is small. Mann (1986, p. 22) estimated inflow to the aquifer at the INL from underlying rocks to be 20 ft3/s on the basis of data from a 10,365-ft-deep test hole drilled at the INL. These estimates were based on the steepest vertical gradients measured at the INL. Extrapolation of that estimate to the 1,940 mi2 of the model area indicates an inflow of about 44 ft3/s or a flux of about 0.02 (ft3/s)/mi2—or about half the flux estimate for precipitation recharge, 0.04 (ft3/s)/mi2.
All inflows across the northwest mountain-front boundary, conceptualized as nonuniform and constant, consist of underflow from alluvial aquifers in the tributary valleys of the Big Lost River (367 ft3/s), Little Lost River (226 ft3/s), and Birch Creek (102 ft3/s). Because no estimates of underflow from the Paleozoic carbonate rocks along the mountain fronts were available, and the contribution from mountain fronts to underflow across the boundary was assumed to be small, the mountain fronts were treated as no-flow sections along the northwest mountain-front boundary. Hydraulic gradients abruptly increase near the mouths of the tributary valleys along the northwest mountain-front boundary, and calculated gradients at the mouths of the Big Lost River, Little Lost River, and Birch Creek are 150, 200, and 120 ft/mi, respectively. These gradients are relatively constant because the variability of the gradients is small compared to the total change in gradients. In a study of the water resources in the Big Lost River basin, Crosthwaite and others (1970, p. 97–98) concluded that the underflow from the Big Lost River basin downstream of Arco is nearly constant each year. Mundorff and others (1964, p. 185) calculated basin yields for the Big Lost River (456 ft3/s), Little Lost River (205 ft3/s), and Birch Creek (110 ft3/s). Kjelstrom (1986) calculated mean annual basin yields for the Big Lost River, Little Lost River, and Birch Creek for water years from 1934 to 1980 of 462, 226, and 102 ft3/s, respectively. These values included contributions from streamflow, alluvial-valley underflow, and underflow from the rocks at the mountain fronts. Subtracting the 95 ft3/s estimated average annual streamflow for the Big Lost River (Brennan and others, 2004, p. 196) from Kjelstrom’s (1986) basin-yield estimate, the underflow from the Big Lost River is 367 ft3/s.
Underflow across the northeast boundary from the regional aquifer, conceptualized as nonuniform and constant, was estimated as 1,225 ft3/s for the model area using Darcy’s Law, a hydraulic conductivity of 140 ft/d (a small value to reflect the presence of the less transmissive sediments), an average hydraulic gradient of 36 ft/mi, a boundary length of 35 mi, and a saturated thickness of 600 ft. The estimate is about 20 percent smaller than the estimate of underflow across the boundary (1,500 ft3/s) summed from a flow net of the eastern SRP aquifer by Mundorff and others (1964, pl. 4), and was used for the conceptual model because it was derived from (1) more numerous and better distributed boreholes in the area, (2) a longer record of water-level measurements, and (3) more numerous measurements of transmissivity and hydraulic conductivity for the area (Ackerman, 1991, table 2; Spinazola, 1994, p. 41 and fig. 35; Bartholomay and others, 1997, table 3; Anderson and others, 1999, table 2). The average saturated thickness along the northeast boundary of the model area was estimated to be about 600 ft on the basis of Whitehead’s (1986, sheet 2) interpretation of the thickness of basalt and sediment, digital elevation models, and the 1980 potentiometric surface (Lindholm and others, 1988). Garabedian (1986, table 13) estimated transmissivity near the northeast boundary of 60,000 ft2/d and Norvitch and others (1969, fig. 12) estimated 700,000 ft2/d. Using the saturated thickness of 600 ft, the corresponding hydraulic conductivity ranges from 100 to about 1,000 ft/d. Using the saturated thickness of 600 ft and transmissivity values at the northeast boundary of 100,000 ft2/d from Mundorff and others (1964, pl. 6) and less than 20,000 ft2/d from Newton (1978, table 16), the corresponding hydraulic conductivities are less than 200 ft/d and about 30 ft/d, respectively. Garabedian (1992, pl. 6) estimated a range of hydraulic conductivity near the boundary of from 55 to 130 ft/d. Anderson and others (1999, table 3) reported geometric means and arithmetic means for hydraulic conductivity of 140 ft/d and 1,500 ft/d, respectively.
The southeast-flowline boundary is treated as a noflow boundary. The boundary is represented as a flowline projected vertically to depth, and consequently flow across this boundary, conceptualized as uniform and constant, is presumed to be zero. Water-table maps constructed for the eastern SRP aquifer for 1928–30, 1956–58, and 1980 (Stearns and others, 1938; Mundorff and others, 1964; and Lindholm and others, 1988) as shown by Garabedian (1992, pl. 4) indicate that changes in water levels near the INL are spatially and temporally uniform. Consequently, flow directions along most of the southeast boundary of the model area are relatively stable, and flowlines constructed for this area should be temporally constant.
Outflow across the southwest boundary (fig. 17) was represented as nonuniform and variable underflow to the regional aquifer and was estimated at 2,037 ft3/s. This value was estimated for the model area using Darcy’s Law and data from previous studies, and is similar to the value, 2,000 ft3/s, summed for underflow across this boundary from the flow net by Mundorff and others (1964, pl. 4).
The true physical length and shape of this boundary are not well known and the water-table contours and hydraulic gradients in the area are not well defined because head definition in this part of the aquifer was limited to water-level measurements in only six wells that are from 3 to 10 mi from the boundary. The southwest boundary was estimated to be about 24-mi long; however, water-table contours indicate that flow may not cross some sections of the boundary (Lindholm and others, 1988). Whitehead (1986, sheet 2) estimated the average saturated thickness of the aquifer in the southwestern part of the model area to be about 2,000 ft. Estimates of transmissivity in this area range from about 1,700,000 ft2/d (Norvitch and others, 1969, p. 37) to about 2,300,000 ft2/d (Mundorff and others, 1964, pl. 6). Assuming an aquifer thickness of 2,000 ft the corresponding hydraulic conductivity ranges from 850 to 1,150 ft/d. Using Darcy’s Law, the smallest (850 ft/d) and largest (1,150 ft/d) estimates of hydraulic conductivity, a boundary length of about 22 mi (across which flow occurs), a hydraulic gradient of 4 ft/mi (fig. 3), and a saturated thickness of 2,000 ft, underflow across the southwest boundary ranges from 1,731 to 2,343 ft3/s. The average of these values, 2,037 ft3/s, was used in the water budget for the conceptual model.
A ground-water budget was derived for the conceptual model area for calendar year 1980 from estimates of inflows and outflows across the model boundaries. A budget residual of about 7 percent between inflow and outflow results from the uncertainty in the estimates of flow across each boundary using data from different sources. These ground-water budget components were compared with components calculated for the conceptual model area using output from the RASA model (table 6) (Garabedian, 1992). Even though the conceptual model area covers only about 18 percent of the RASA model area (fig. 5), the flows calculated for the conceptual model from the RASA model were considered reasonably accurate because the RASA model was constrained by measured outflows in the Thousand Springs area (fig. 3).
Of the inflow components of the ground-water budget for the model area, about 8 percent enters at the water table, releases from storage account for about 4 percent, upward flow across the base of the aquifer contributes about 2 percent, and about 86 percent enters as underflow along the northwest mountain-front and northeast boundaries (table 6). Of the 8 percent of the budget that enters the model area at the water table, 3 percent is by recharge from precipitation, 4 percent is from streamflow-infiltration recharge along the Big Lost River, and 1 percent is from infiltration of irrigation and industrial wastewater return flows. Underflow along the northwest mountain-front and northeast boundaries contribute 31 percent and 55 percent, respectively, of the total inflow into the model area. Flow across the southeast-flowline boundary, approximated by a pathline projected from depth to the water table (Ackerman, 1995), was designated as zero because it was represented as a flowline or noflow boundary. At the scale of the conceptual model area, this noflow boundary is an idealized approximation, but some local flow probably occurs across this boundary.
Of the outflow components of the ground-water budget for the conceptual model area, about 2 percent is from ground‑water withdrawals and about 98 percent is underflow across the southwest boundary to adjacent parts of the aquifer. The ground-water withdrawals are the amount of ground water withdrawn for irrigation and industrial uses.
For comparison purposes, a ground-water budget for the conceptual model area, which lies entirely in the RASA model area (fig. 5), was calculated using input and output from the transient three-dimensional RASA model simulations (table 6) (Garabedian, 1992). The RASA-derived budget was calculated for the period 1976–80, a time when hydrologic conditions in the eastern SRP aquifer were relatively stable. Flow across each conceptual model boundary was calculated by tabulating flow values for each regional model cell (4 mi on a side) that most closely approximated conceptual model boundaries.
Total inflow and total outflow from the RASA-derived ground-water budget exceeded the ground-water budget estimates for the conceptual model by 14 and 23 percent, respectively (table 6). These percentage differences are relatively large. Unlike the conceptual model budget, the RASA budget was derived from a transient model simulation that, by design, precluded a budget residual because differences between inflow and outflow within every grid cell were accounted for by changes in storage. Residual in the conceptual model budget, 157 ft3/s (table 6), represents the equivalent of 6 percent of the inflow and outflow components of the RASA-derived budget. To some extent, the 157 ft3/s residual is a measure of the cumulative uncertainty of the inflow and outflow estimates and may account for some of the differences in the inflow (311 ft3/s) and outflow (469 ft3/s) estimates of the two water budgets.
Some of the differences in the individual inflow and outflow estimates of the two water budgets can be attributed to a combination of (1) coarse discretization of the RASA model grid, (2) interpolation errors resulting from inflow and outflow estimates that are assumed to be uniformly distributed within the RASA model grid cells and that accumulate where the grid cell faces approximating the conceptual model boundaries are offset from the grid cell faces of the RASA model (fig. 16), (3) noncompensating summation errors, particularly for estimates along the southeast boundary, which is oriented parallel to the direction of regional flow and is therefore particularly susceptible to this type of error (fig. 16), (4) hydrogeologic judgment, and (5) accounting methodology.
The 79 ft3/s difference between the conceptual-model and RASA-derived estimates of outflow from ground-water withdrawals reflects accounting limitations owing to coarse discretization of the model grid. Most ground-water withdrawals for irrigation are in the Mud Lake area, north of the northeast boundary of the conceptual model. The grid cells in the RASA model that approximate the alignment of the northeast boundary of the conceptual model encompass more wells than are actually present in the conceptual model area (fig. 16) as summed from Spinazola’s (1994) and McAndrew’s (U.S. Bureau of Reclamation, written commun., 1996) data.
Interpolation errors were associated with estimates of underflow across the south corner of the southwest boundary of the conceptual model. Estimates of underflow across this boundary, for which data are very limited, are perhaps the least reliable of all the water-budget components. These estimates were based on application of Darcy’s Law. To account for uncertainties in hydraulic conductivities, hydraulic gradients, and saturated thickness of the aquifer in the southwestern part of the model area, various estimates of these factors were used to compute underflow estimates. Underflow estimates ranged from 1,731 to 2,343 ft3/s; the average of this range, 2,037 ft3/s, was used for the conceptual model. This range represents a difference of 612 ft3/s, or about 25 percent of the 2,427 ft3/s RASA-derived estimate for underflow across the southwest boundary. The distribution of flow along the 22-mi length of the southwest boundary of the conceptual model, for each vertical stack of cell faces in the three-layer RASA model, ranges from a minimum of 1.2 ft3/s at the north corner of this boundary to a maximum of 606 ft3/s at the south corner, and averaged 404 ft3/s. Because the alignment of the southwest boundary of the conceptual model is coincident with a grid line in the RASA model (fig. 16), interpolation errors were limited to the south corner of the boundary, where interpolation of the 606 ft3/s outflow across the vertical stack of cell faces at this corner of the boundary can introduce a substantial difference in outflow estimates.
Noncompensating summation errors were the likely source of budget differences for the southeast-flowline boundary, which was treated as a noflow boundary in the conceptual model. Inflows across the southeast-flowline boundary of the conceptual model were based on the summation of interpolated flows into and out of a stair‑stepped array of 19 three-layer stacks of cells consisting of 63 cell faces in the RASA model (Garabedian, 1992) that approximate the southeast boundary of the conceptual model (fig. 16). Along this boundary, 10 three-layer vertical cell stacks produce a net outflow and 11 produce a net inflow. Flow into or out of a single three-layer vertical stack of cell faces along this boundary ranges from a minimum of 9 ft3/s to a maximum of 594 ft3/s and averages 125 ft3/s. Summation of inflows and outflows across this boundary were susceptible to both interpolation errors and summation errors that probably are not fully compensating, and these kinds of errors could have resulted in the 194 ft3/s difference between the conceptual-model inflow estimates and the RASA-derived inflow estimates (table 6).
Precipitation-recharge estimates for the RASA-derived budget were 87 ft3/s larger than those for the conceptual model budget, but most of this difference represents a 59-percent allocation (92 ft3/s) of the total precipitation recharge (157 ft3/s) in the RASA-derived budget to the southwest corner of the conceptual model, an area that represents 18 percent (352 mi2) of the conceptual model area (fig. 16). In the RASA model, this allocation represents a precipitation-recharge flux of 0.26 (ft3/s)/mi2. This flux is more than six times larger than estimates for the remainder of the conceptual model area, 0.04 (ft3/s)/ mi2. The larger precipitation-recharge estimate for this area in the RASA model was attributed to the presence of bare basalt outcrops and the assumed effects of higher altitude on precipitation. However, there are few data to justify this large estimate.
Other noteworthy differences in the two ground-water budgets are attributed to differences in accounting methodology. These include (1) an irrigation infiltration component of 24 ft3/s in the conceptual model budget that was accounted for in the surface-water irrigation component of the RASA-derived budget, (2) flow across the base of the aquifer that was assumed to be zero in the RASA-derived budget but was estimated to be as large as 44 ft3/s in the conceptual model budget, and (3) northwest mountain-front boundary underflow of 48 ft3/s in the RASA-derived budget that was not accounted for separately in the conceptual model budget.
Inflows from the northwest mountain-front boundary, the northeast boundary, and streamflow infiltration are similar in the two budgets. Combined inflows of 2,054 ft3/s for these components in the RASA-derived budget compare favorably with combined inflows of 2,015 ft3/s for these components in the conceptual model budget.
Differences in the two water budgets emphasize the probable magnitude of the uncertainty associated with current water-budget estimates. This uncertainty is approximately ±7 percent of the average value of the total inflow (2,395 ft3/s) and ±10 percent of the average value of the total outflow (2,317 ft3/s) of the two budgets. These averages do not include any adjustments for the residual flow shown in table 6.
Water-table maps, based primarily on water levels from the upper 200 ft of the of the eastern SRP aquifer, indicate that the direction of regional ground-water flow is from northeast to southwest (fig. 3) (Garabedian, 1992, fig. 26). Water-table maps for the central part of the INL, however, indicate ground water in this area flows southward and southeastward. Locally, flow directions in the INL vary from southeast to southwest in response to episodic recharge from streamflow-infiltration. Flow increases progressively in a direction downgradient of the northeast boundary because of mountain-front and tributary-valley underflows along the northwest boundary and precipitation recharge and streamflow-infiltration recharge across the water-table boundary. Ground-water flows through all three hydrogeologic units beneath the INL. In the northern part of the INL and south of the INL, the younger rocks (hydrogeologic units 1 and 2) are either absent or are above the water table and all flow takes place through the slightly altered basalts composing the intermediate-age rocks (hydrogeologic unit 3).
Horizontal hydraulic gradients and the direction and velocity of horizontal ground-water movement have been inferred from water-table maps and interpretations of water-chemistry data and are primarily representative of the upper 200 ft of the aquifer in the model area. The earliest water-table maps for the eastern SRP aquifer were compiled during 1928–30 (Stearns and others, 1938) and, over the past 50 years, tens of thousands of aquifer water-level measurements were collected by the USGS. In 2004, for example, about 670 water levels were measured at 172 aquifer wells and 36 perched-water wells in the model area. Water-chemistry data were routinely collected by the USGS in the model area since 1950. In 2004, for example, water samples were collected at 133 aquifer wells, 26 perched-water wells, and 7 surface-water sites. Water-chemistry data referred to in this report include (1) major and trace elements, (2) stable isotope ratios, (3) isotopes resulting from water-rock interactions, (4) atmospheric tracers, and (5) concentrations or ratios of contaminants discharged at facilities.
Water moves horizontally in the aquifer principally through porous interflow zones between basalt flows rather than in the massive interiors of basalt flows (Whitehead, 1992, p. 26). Because interflow zones are more prevalent in the densely fractured basalt of hydrogeologic unit 1 than in the massive, less densely fractured basalt of hydrogeologic unit 2, the amount of water moving through hydrogeologic unit 1 is probably much larger than that moving through hydrogeologic unit 2.
Ground water moves vertically in the aquifer principally through fracture systems (Mann, 1986, p. 21). Water also moves vertically through fine-grained sediment and massive basalts, but vertical movement probably is impeded substantially by these relatively horizontal, impermeable layers. Few data were available to define the vertical hydraulic gradients and the vertical movement of ground water in the aquifer. The data that were available were (1) measurements of hydraulic head at several depths during drilling and testing of the 10,365-ft deep INEL 1 test hole (fig. 1), (2) water-level measurements from dedicated piezometers located at various depths in wells USGS 30 and Hwy 1 (fig. 1), and (3) intra-borehole flow measurements of upward or downward ground-water flow in 16 wells located near TAN, the Naval Reactors Facility (NRF), the TRA, and the INTEC (fig. 1).
Horizontal hydraulic gradients vary spatially across the model area in response to local changes in hydraulic conductivity, aquifer thickness, and recharge. These water-table gradients were defined by measurements of water levels in 201 aquifer wells in the model area, of which 175 wells are in the INL. A few wells have multiple completions, and of the 207 well completions 158 penetrate only the upper 200 ft of the aquifer. At a regional scale, water-table gradients have remained relatively stable during the past 70 years (see water-table maps in Stearns and others, 1938; Garabedian, 1992, plate 4; and Bartholomay and others, 2000, fig. 9). Locally, however, the gradients may fluctuate annually in response to streamflow-inflitration and irrigation recharge and cyclically over 3- to 8-year periods in response to climate variation and its effect on the quantity of streamflow-infiltration recharge.
Water-table gradients range from about 1 to 60 ft/mi (fig. 17) in the model area and from about 1 to 8 ft/mi beneath the INL (fig. 12), and precise definition of flow direction beneath the INL is difficult owing to the small gradient. Gradients near the northeast boundary of the model area range from 27 to 60 ft/mi (fig. 16) (Lindholm and others, 1988). These large gradients probably are caused by a decrease in hydraulic conductivity associated with fine-grained sediments (Nace and others, 1959, p. 151; Mundorff and others, 1964, p. 133; Lindholm and others, 1988). The gradient in the northern part of the INL is poorly defined, but was estimated to be about 1 ft/mi near TAN (figs. 1 and 12) (Anderson and Liszewski, 1997, p. 20; Sorenson and others, 2000, p. 340). In much of the central part of the INL the gradient is about 6 to 8 ft/mi, and this larger gradient probably reflects the larger amount of sediment and the resulting smaller hydraulic conductivity in this area (fig. 12) (Anderson and Liszewski, 1997, p. 20-21). Water-level contours (fig. 12) indicate a gradient of about 2 to 4 ft/mi in most of the southern and eastern parts of the INL, except in the area north and east of Middle and East Buttes, where gradients are about 4 to 6 ft/mi. The smaller gradient in the southern and eastern parts of the INL may reflect increased transmissivity resulting from less sediment and greater aquifer thickness in this area (figs. 12 and 14), and the larger gradient north and east of Middle and East Buttes probably results from the smaller hydraulic conductivity associated with the silicic rocks in this area. Gradients southwest of the INL range from about 4 to 30 ft/mi (fig. 17) (Lindholm and others, 1988). The larger gradients may result from structural uplift in this area, perhaps related to differential subsidence or faulting, or the emplacement of laccoliths and domes within the stratigraphic section (Anderson and others, 1997, p. 5 and fig. 3), causing the relatively impermeable basalts of hydrogeologic unit 2 to intersect the water table (figs. 9 and 10). Alternatively, Lindholm and others (1988) noted that this large gradient is immediately upgradient of, and generally parallel to, a zone of smaller transmissivity. This zone of smaller transmissivity may be attributed to the healing of fractures or the presence of numerous, closely spaced dikes that may be related to the nearby and downgradient Great Rift volcanic rift zone (Lindholm and others, 1988; Anderson and others, 1999, p. 26). In either case, larger gradients would be necessary to move water through the low-permeability basalts of hydrogeologic unit 2 or the low-permeability rocks associated with the rift zone.
Although ground water moves southwestward regionally, ground water in the central part of the INL was inferred to flow southward and southeastward on the basis of water-table contours (fig. 12), hydraulic-gradient analyses from water-level triangulations (Michael Rohe, Bechtel, Babcock, and Wilcox, Idaho, written commun., 2000), and interpretation of water-chemistry data (Sorenson and others, 1996, figs. 2‑23; Johnson and others, 2000, p. 873; Luo and others, 2000, fig. 9b; Busenberg and others, 2001, fig. 25). The southward and southeastward flow directions may result from the flow directions of tributary valley and regional underflow and directional changes in aquifer transmissivity.
The variation in flow direction may be attributable to the two principal sources of water to the aquifer beneath the INL: (1) regional underflow from the northeast (55 percent of total estimated inflow for 1980; table 6) and (2) tributary-valley underflow (Big Lost River, Little Lost River, and Birch Creek valleys) from the northwest (31 percent of total estimated inflow for 1980; table 6). Underflow from the Little Lost River and Birch Creek valleys (15 percent of total estimated inflow for 1980; table 6) enters the eastern SRP aquifer in the north-central and northern parts of the INL (fig. 17), respectively. These valleys trend southeastward and, given the large hydraulic gradients at the mouths of these valleys, underflow entering the eastern SRP aquifer from these valleys would be expected to flow in the same direction. When the tributary valley underflow converges with regional underflow, which moves in a southwesterly direction, the combined gradients driving these flows may produce flow that moves in a southerly direction.
Evidence that tributary-valley and regional underflow converge in the central part of the INL is provided by the chemical signatures of ground water in this area. Regional-source and tributary-valley source ground water are distinguishable beneath the INL because these ground waters originate in silicic and carbonate aquifers, respectively, and contain distinctive chemical signatures (Olmstead, 1962, p. 37‑38; Robertson and others, 1974, p. 42-55). Concentrations of major ions (potassium, sodium, calcium, magnesium, fluoride, bicarbonate), silica, trace elements (lithium, boron, strontium), and stable isotope ratios (delta oxygen-18, delta carbon-13) in waters from the eastern SRP aquifer (Olmstead, 1962, p. 37-38; Robertson and others, 1974, p. 42-55; Busenberg and others, 2001, figs. 3-5 and 35‑36), indicate that ground water beneath the southeastern part of the INL retains the chemical signature of regional ground water and ground water beneath the northwestern part of the INL retains the chemical signature of ground water from the tributary valleys. For example, lithium (Li) concentrations beneath the southeastern part of the INL are greater than 5 µg/L, similar to those in regional ground water, and beneath the northwestern part of the INL they are less than 5 µg/L, similar to those in tributary-valley ground water (fig. 19). The lithium line in figure 19 separates the area of Li concentrations in the aquifer attributed to regional ground water (greater than 5 µg/L) from that attributed to tributary-valley ground water (less than 5 µg/L), and the line may indicate the area where regional and tributary-valley ground water converge.
Ground water also was interpreted to flow southward through the central and northern parts of the INL by identifying water-rock interactions that produced chemically‑evolved isotope concentrations (uranium and thorium decay-series disequilibria: uranium, thorium, and radium isotopes) (Luo and others, 2000, fig. 9b) and distinctive stable isotope ratios (strontium-87/strontium-86) (Johnson and others, 2000, fig. 2a). Fast-flow zones were interpreted to extend south from the mouths of the Little Lost River and Birch Creek and slow-flow zones were interpreted to be present in the western and central parts of the INL. These fast- and slow-flow zones, however, may not be consistent with ground-water flow velocities calculated by Busenberg and others (1993, 2001), as shown on figure 20, on the basis of model ages of water using atmospheric tracers (3H/3He, chlorofluorocarbon) and contaminant ratios (3H/3He). The model ages indicated slower flow (2 to 8 ft/d) between the mouth of the Little Lost River and the INTEC and faster flow (10 to 14 ft/d) in the central part of the INL (fig. 20) (Busenberg and others, 1993, p. 30; Busenberg and others, 2001, fig. 25). Busenberg and others (2001, p. 54) state that the results of their study are consistent with the interpretation of a fast-flow zone in the central part of the INL by Luo and others (2000, fig. 9b) and Johnson and others (2000, fig. 2a). Although the results of Busenberg and others (2001, fig. 25) support a conclusion that fast- and slow-flow zones exist in the aquifer beneath the INL, the fast-flow zone in the central part of the INL indicated by Busenberg and others (2001, fig. 25) is not located within the fast-flow zones indicated by Luo and others (2000, fig. 9b) and Johnson and others (2000, fig. 2a).
The transition from southerly flow in the central part of the INL to regional, southwesterly flow also may reflect differences in aquifer transmissivity beneath the INL. The aquifer is thinner near the mountain fronts, thickens to the southeast, and is thickest along the southeast model boundary (fig. 14). These aquifer thicknesses and the distribution of hydrogeologic units and hydraulic properties described in the section, “Hydrogeologic Framework and Hydraulic Properties,” (figs. 9 and 12, table 2), indicate that the transmissivity of the aquifer probably is (1) smaller in the northern and southwestern parts of the INL, (2) larger in the central and eastern parts of the INL, and (3) much smaller in the northwestern than in the southeastern part of the INL. This smaller-to-larger northwest-to-southeast difference in aquifer transmissivity indicates that there is less regional underflow through the northwestern than through the southeastern part of the aquifer. The lesser volume of regional underflow in the northwestern part of the INL probably allows the volume and direction of tributary-valley underflow to exert a relatively large influence on shallow flow directions in the northern and central part of the INL.
Flow directions beneath the southwestern part of the INL vary temporally because of gradient reversals produced by episodic streamflow-infiltration recharge from the Big Lost River (Barraclough and others, 1976, p. 52-56). Flow in the Big Lost River is usually largest in the spring and early summer when snow melts in the high-altitude river basin and, during years with sufficient precipitation, runoff reaches the central and southwestern parts of the INL where it infiltrates and recharges the aquifer. During wet years, flow in the Big Lost River in the central part of the INL was more than three times (190 ft3/s) the average annual streamflow (55 ft3/s) and flow diverted to the spreading areas was more than seven times (379 ft3/s) the average annual streamflow (48 ft3/s) (table 5; streamflow-gaging stations 13132520 and 13132513, respectively). This large amount of cold surface-water recharge probably acts as a persistent heat sink in the aquifer and could have depressed ground-water temperatures in these areas (fig. 21).
Pulses of streamflow-infiltration recharge have created large localized water-level changes and water-table mounds (fig. 22) in the low-gradient area beneath the central INL (fig. 17). For example, in the early 1980s larger than normal streamflow infiltrated and recharged the aquifer during a wet climate cycle. The water table rose unevenly beneath the INL, and water-table mounds formed near the Big Lost River spreading areas and in the area north of the NRF. The water table fell unevenly during a dry cycle in the late 1980s and early 1990s, when there was no streamflow recharge, and the water-table mounds dissipated. Water-table mounds were less pronounced during a wet cycle in the late 1990s when the water table rose moderately.
Nonuniform water-level changes are typical beneath the INL, with the largest changes occurring near or beneath areas of episodic recharge. Hydrographs of water levels in well USGS 9 near the spreading areas and well USGS 12 north of the NRF show water-level rises from 13 to 16 ft during and following the wet climate cycle in the early 1980s, and declines from 16 to 22 ft during the dry climate cycle of the late 1980s and early 1990s (fig. 23). The water level at well USGS 9 reached its highest peak of the period 1948–2003 in 1984 (fig. 23), the second year of a 2-year period of record-high diversions to the spreading areas (fig. 18, gaging station 13132513). The water level in well USGS 9 dropped rapidly after diversions to the spreading areas decreased during 1985. In contrast, the high water levels in well USGS 12 that persisted from 1985 through 1987 (fig. 23) reflect the more uniform flow in the Big Lost River downstream of the INL diversion dam (fig. 18, gaging station 13132520) between 1982 and 1986. During the wet climate cycle of the early 1980s, the rise in water levels at wells USGS 25 and Arbor Test, in the northern and southeastern parts of the INL (fig. 1), was about 7 and 11 ft, respectively (fig. 23). The smaller increase in water levels at these wells in response to the wet climate cycle probably reflects their greater distance from the localized pulses of streamflow-infiltration recharge. A similar but smaller rise in water levels at these four wells was recorded during the wet climate cycle of the late 1990s (fig. 23).
Ground water in and near a water-table mound would be expected to flow radially away from the mound and, at the INL, water east of the mound would be expected to flow southward and southeastward around the mound. An example of how local flow directions may vary, and may even reverse direction owing to water-table mounds was presented by Barraclough and others (1976, p. 52-56 and fig. 33). On the basis of water-level measurements in the vicinity of the RWMC during 1972, Barraclough and others (1976, p. 52‑56 and fig. 33) inferred that ground water flowed away from and around a water-table mound near the Big Lost River spreading areas (fig. 1). They noted that the northeast direction of flow indicated by these data was counter to the prevailing southwestern regional flow direction, and suggested that streamflow-infiltration recharge from the Big Lost River spreading areas appeared to be the predominant influence on this anomalous flow direction.
The variation in the direction of ground-water flow in response to pulses of streamflow-infiltration recharge may be large, perhaps tens of degrees. Simulated flow directions in a flow-and-transport model (Goode and Konikow, 1990b, p. 420, fig. 2) varied by as much as 20 degrees, between N. 70° E. to N. 90° E., about 700 ft downgradient of the disposal well (CPP 3) at the INTEC in response to variable streamflow-infiltration recharge from the Big Lost River. Larger, unreported, variations in flow directions may have been present at other locations in the model. Flow directions west of INTEC, near the Big Lost River, were interpreted to range from about N. 80° E. to N. 280° W. during 1976–1998 based on a study evaluating the temporal fluctuation of hydraulic gradients at the INL (Michael Rohe, Bechtel, Babcock, and Wilcox, Idaho, written commun., 2000).
The estimated average linear ground-water velocities in the model area range from 2 to 20 ft/d (fig. 20) and were determined from concentrations or ratios of atmospheric tracers (3H/3He, chlorofluorocarbons) and long-term monitoring of contaminant movement in the aquifer (3H, 129I, 36Cl, 3H/3He). Velocities determined using atmospheric tracers were estimated from the 3H/3He- or chlorofluorocarbon-model age of the water and the probable location of recharge (Busenberg and others, 1993, p. 29; Busenberg and others, 2001, p. 41, fig. 25). Velocities determined using concentrations or ratios of contaminants discharged at facilities were calculated based on assumed first-arrival times of contaminants at wells downgradient of known contaminant input locations (Barraclough and others, 1981, p. 59; Pittman and others, 1988, p. 51; Mann and Beasley, 1994, p. 24; Cecil and others, 2000, p. 686). Velocities were estimated to be 2 ft/d near TAN (Busenberg and others, 1993, p. 30; Busenberg and others, 2001, fig. 25); 15 to 18 ft/d between Mud Lake and Atomic City (Busenberg and others, 1993, p. 30); 2 to 14 ft/d in the area north and east of the TRA and southeast of Howe (Busenberg and others, 1993, p. 30; Busenberg and others, 2001, fig. 25); 4 to 20 ft/d south of the TRA and INTEC (Barraclough and others, 1981, p. 59; Pittman and others, 1988, p. 51; Mann and Beasley, 1994, p. 24; Cecil and others, 2000, p. 686; Busenberg and others, 2001, fig. 25); and 3 to 4 ft/d south of the Big Lost River and near the western boundary of the INL (Busenberg and others, 1993, p. 30; Busenberg and others, 2001, fig. 25) (fig. 20).
These estimated average linear ground-water velocities generally correlate with the ranges and spatial distributions of hydraulic conductivity values estimated previously in the section, “Hydrogeologic Units” (table 2). High ground-water velocities were estimated for areas where the highly conductive, densely fractured basalt (hydrogeologic unit 1) is present at the water table, and lower velocities were estimated for areas where less conductive, less densely fractured basalt (hydrogeologic unit 2) and intermediate-age rocks (hydrogeologic unit 3) are present at the water table (fig. 9). Using the range of velocities presented above, an average hydraulic gradient at the INL of 4 ft/mi, a range of effective porosities of 0.1 to 0.2 for fractured basalt and 0.05 to 0.15 for massive basalt and intermediate-age rocks, calculated horizontal hydraulic conductivities range from 300 to 5,000 ft/d, 200 to 800 ft/d, and 100 to 400 ft/d for the densely fractured basalt (hydrogeologic unit 1), less-densely fractured basalt (hydrogeologic unit 2), and intermediate-age rocks (hydrogeologic unit 3), respectively.
Hydraulic conductivities derived from velocity estimates for the area of ground-water contamination south of the TRA and INTEC range from 500 to 5,000 ft/day for assumed effective porosities of 0.1 and 0.2. These estimates are generally consistent with results of aquifer testes conducted in hydrogeologic unit 1. The ground-water velocity in hydrogeologic unit 1 in the range from 4 to 20 ft/d probably reflects preferential flow along the many interflow zones of the thin, densely fractured basalts composing the uppermost hydrogeologic unit.
High wastewater-discharge rates at disposal wells may produce increased average linear ground-water velocities downgradient of the wells. Evidence of an increase in ground‑water velocity in the area between the disposal well (CPP 3) at the INTEC and two wells (USGS 40 and USGS 41, in fig. 1) less than 1,000 ft to the southwest was provided in two tracer studies conducted one year apart (Jones, 1963). During the first tracer study, the rate of wastewater discharged at CPP 3 was gradually increased from less than 100 gal/min to about 700 gal/min. During the second tracer study, the discharge rate remained steady at about 970 gal/min (Jones, 1963, p. 228). The ground-water velocity between the two wells calculated in the second study was 2 to 4 times higher than that calculated in the first study (Jones, 1963, p. 228). The most probable explanation for this increase in velocity was the higher wastewater-discharge rate at the disposal well during the second study.
The small number of deep wells and the few measurements of vertical hydraulic head and upward or downward ground-water movement provided limited information about the vertical gradients and movement of ground water in the model area, but did demonstrate that the direction of vertical movement is spatially variable. The information available regarding vertical hydraulic conductivity and head indicates that hydraulic conductivity decreases (Mann, 1986, p. 6-14) and hydraulic head increases (Jones, 1961, p. 41; Morin and others, 1993, p. 16-19; Mann, 1986, p. 14) with depth in the southwestern part of the INL. However, in the upper part of the aquifer, near TAN, the NRF, the TRA, and the INTEC, measurements of ground-water movement in wells indicates that water moves both upward and downward (Morris and others, 1964, p. 40-42; Morris and others, 1965, p. 42-44; Barraclough, Teasdale, and Jensen, 1967, p. 94-98; Morin and others, 1993, table 1).
Water-level measurements2 at test hole INEL 1 (fig. 1) show that the hydraulic head at depth (4,210–10,365 ft below land surface) is about 115 ft higher than the head for the upper 200 ft (395–595 ft below land surface) of the aquifer (Mann, 1986, p. 14). Water-level measurements in piezometers at wells USGS 30 and Hwy 1, in the eastern part of the INL and east of the southeast boundary of the model area (fig. 1), respectively, also show head increasing with depth. At well USGS 30, head increases of 2 and 11 ft were measured at depths of 335 and 545 ft below the water table, respectively, and at well Hwy 1, head increases of 0.4 and 0.3 ft were measured at depths of 100 and 425 ft below the water table, respectively. The increased head with depth at wells USGS 30 and Hwy 1 may be related to the steep hydraulic gradient along the northeast boundary of the model area (fig. 3). The steep gradient coincides with the interfingering of basalt with less transmissive layers of sediment (discussed in the section “Artificial Boundaries”). These sediment layers may act as confining units and, if the deeper sediment layers extend further southwestward than the shallower layers, the heads in the northeastern part of the model area could be larger beneath the deeper sediment layers than in the overlying rocks.
Variation in the direction of vertical ground-water flow in the upper 300 ft of the aquifer is confirmed by measurements of intra-borehole flow at 16 wells. Downward flow was measured at one well3 (USGS 4) southeast of TAN, one well (Site 17) north of the NRF, two wells (MTR Test, TRA disposal) at the TRA, and one well (USGS 49) at the INTEC; upward flow was measured at two wells (USGS 28, USGS 31) southeast of TAN and at five wells (USGS 42, USGS 44, USGS 45, USGS 46, USGS 47) at the INTEC; and upward and downward flow was measured at four wells (USGS 48, USGS 51, USGS 52, USGS 59) at the INTEC (Morris and others, 1964, p. 40-42; Morris and others, 1965, p. 42-44; Barraclough, Teasdale, and Jensen, 1967, p. 94-98; Morin and others, 1993, table 1).
2 Hydrologic data on the World Wide Web may be accessed at http://idaho.usgs.gov/
3 Wells USGS 4, USGS 28, USGS 31, and Site 17 are shown in figure 1. The other wells discussed in this sentence are located inside or near the indicated facilities and are not shown on map. The facility locations are shown in figure 1.
Ground-water flow through the model area is controlled by the structural and stratigraphic framework of the aquifer and the distribution and timing of hydrologic inputs to and outputs from the model area. Although the conceptual model simplifies geologic and hydrologic features it retains features of the system important for modeling contaminant transport.
Structural features in the model area that may control ground-water flow in the aquifer by providing preferential pathways for, and impediments to, flow include (1) rhyolite domes, (2) a sedimentary trough, and (3) volcanic rift zones and vent corridors. Smaller hydraulic conductivities associated with rhyolite domes and rocks in the sedimentary trough decrease flow velocities and may locally alter flow directions where these features are present. Volcanic rift zones and vent corridors may produce both porous, rapid flow zones and vertical, low permeability structures that impede flow.
The conceptual model represents rhyolite domes, areas of subsidence, uplift, and dipping beds, and the sedimentary trough. It does not represent volcanic rift zones or vent corridors. The effects of volcanic rift zones and vent corridors on ground-water movement in the eastern SRP aquifer are undetermined. Stratigraphic controls on ground-water flow in the model area are the horizontal and subhorizontal layers of basalt and sediment in the eastern SRP aquifer. In the conceptual model, three homogeneous and anisotropic hydrogeologic units (fig. 10) represent the stratigraphic layers. Hydraulic conductivity of each of the three hydrogeologic units reflects the aggregate lithology, thickness, and number of basalt flows and sedimentary interbeds in each hydrogeologic unit. Stratigraphic layers produce both large- and small-scale heterogeneity and horizontal-to-vertical anisotropy in the aquifer.
The large-scale heterogeneity and anisotropy in the aquifer are represented by the subhorizontal layering and differences in hydraulic conductivity among adjacent homogeneous and anisotropic hydrogeologic units (interunit). The difference in conductivity is particularly prominent between hydrogeologic units 1 and 2, with the conductivity of hydrogeologic unit 2 estimated to be much smaller than hydrogeologic unit 1. Consequently, hydrogeologic unit 2 acts as an impediment to the downward movement of water and contaminants, and wastewater that infiltrates through the unsaturated zone or is discharged to the aquifer at the INL probably is constrained to move in hydrogeologic unit 1 beneath the INL.
Small-scale heterogeneity and anisotropy in the aquifer, which exists in and between the individual basalt flows and sedimentary interbeds (intraunit) that compose the hydrogeologic units, are not represented in the conceptual model. The intraunit stratigraphy produces (1) preferential flow paths and interflow zones that facilitate rapid horizontal ground-water movement, and (2) hydraulic barriers, in the form of massive interiors of individual basalt flows and sedimentary interbeds, that substantially impede vertical ground-water movement. Although the small-scale heterogeneity and anisotropy are not represented in the conceptual model, they could be indirectly represented in the numerical model through the choice of values for the horizontal and vertical hydraulic conductivities for each of the hydrogeologic units.
Hydrologic controls on ground-water flow in the model area include regional and tributary-valley underflow, precipitation and streamflow-infiltration recharge, and industrial wastewater return flows. All these hydrologic controls are represented in the conceptual model. Regional and tributary-valley underflow and precipitation recharge can be adequately modeled by steady-state numerical simulations of flow; however, streamflow-infiltration recharge from the Big Lost River and the variable amounts of industrial wastewater return flows requires transient numerical simulations.
The predominant northeast-to-southwest flow direction through the model area is primarily controlled by regional underflow across the northeast boundary, and the south-to-southeast flow direction in the central part of the INL is controlled primarily by underflow entering the eastern SRP aquifer from the Little Lost River and Birch Creek valleys. Regional and tributary-valley underflow entering the model area have undergone gradual changes in magnitude since 1890 in response to climate cycles and changing irrigation practices. However, these changes in underflow do not appear to have affected the magnitude and direction of ground-water flow in the central part of the eastern SRP aquifer. Potentiometric maps of the aquifer indicate that the shape of the water table and flow directions in the model area have remained relatively stable since about 1930 (Garabedian, 1992, p. 21).
In contrast to the gradual changes in underflow entering the model area, large fluctuations were measured in precipitation, streamflow-infiltration recharge, and industrial wastewater return flows. Precipitation and streamflow-infiltration recharge fluctuate annually and in response to short-term (3–8 years), wet-dry climate cycles. Recharge from precipitation in the model area is areally diffuse (about 0.04 (ft3/s)/mi2), and little or no streamflow from the Little Lost River or Birch Creek reaches the INL, so these sources of recharge do not significantly impact water levels or flow directions in the model area. However, because streamflow infiltration from the Big Lost River is spatially focused, fluctuates considerably in response to short-term climate cycles, and can significantly alter water levels and flow directions locally in the model area, it has important consequences for the movement of water and contaminants. Industrial wastewater return flows in the model area have varied in volume and location over time. These flows are small relative to the volume of water flowing through the model area; however, because they occur near contaminated ground water they affect the movement of water and contaminants.
A semi-quantitative distribution of ground-water flow paths is shown on a vertically-exaggerated hydrogeologic section of the model area (fig. 24). Flow paths drawn on the section were based on flow paths generated with a version of the software package FLOWNET (Fetter, 1994, p. 314). Inflows were based on steady-state conditions. Input for the flow net included (1) a uniform, two-dimensional vertical discretization (cells were 4 mi horizontally and 400 ft vertically) of hydrogeologic units 1, 2, and 3 along the line of section shown in figure 9, (2) inflow across the northeast, northwest mountain-front, and water-table boundaries (fig. 17 and table 6, (3) the altitude of the water table along the line of section shown in figure 9, and (4) estimates of the horizontal hydraulic conductivity (Kh) and vertical hydraulic conductivity (Kv) for hydrogeologic units 1, 2, and 3, where
Hydrogeologic unit 1 (younger rocks—thin, densely fractured basalt)
(estimates from Frederick and Johnson, 1996, p. 59).
Hydrogeologic unit 2 (younger rocks—massive, less‑densely fractured basalt)
(estimates from Frederick and Johnson, 1996, p. 59).
Hydrogeologic unit 3 (intermediate-age rocks—slightly altered fractured basalt)
Hydrogeologic unit 3 estimates were assumed to be approximately one order of magnitude smaller than those for hydrogeologic unit 1 (John Welhan, Idaho State Geological Survey, written commun., 1999).
Other assumptions that were used to draw flow paths on the hydrogeologic section included:
Inflows across the water-table boundary were represented on the hydrogeologic section (fig. 24) by superimposing the inflows onto the section. These inflows produced a progressive increase of flow in the aquifer downgradient of the northeast boundary and contributed about 45 percent of the total outflow from the system (table 6).
The hydrogeologic section and semi-quantitative flow paths (fig. 24) indicate that steady-state flow through the aquifer is controlled by the stratigraphic and hydraulic features of the aquifer. The intermediate-age rocks of hydrogeologic unit 3 compose about 75 percent of the aquifer, and most of the water in the aquifer resides in this hydrogeologic unit because of its large volume. Younger rocks of hydrogeologic units 1 and 2 together essentially form a thin layer on top of hydrogeologic unit 3 in the upper and central part of the aquifer. Hydrogeologic unit 1, which is similar in thickness to hydrogeologic unit 2 and much smaller in thickness than hydrogeologic unit 3, has a much larger hydraulic conductivity than that of hydrogeologic units 2 and 3; consequently, hydrogeologic unit 1 transmits a disproportionately large amount of water. The smaller hydraulic conductivity of hydrogeologic unit 2 and the large horizontal-to-vertical anisotropy of the aquifer (Kh is much larger than Kv) generally impedes the downward flow of water from hydrogeologic unit 1 (fig. 24). The pyramid-shaped segment of hydrogeologic unit 2 that extends above the water table in figure 24 is not laterally extensive (fig. 9) and does not significantly impede the horizontal movement of water through hydrogeologic unit 1.
Water moves up and down in the intermediate-age rocks of hydrogeologic unit 3 and, southwest of the INL, water in hydrogeologic unit 1 probably moves down across hydrogeologic unit 2 into hydrogeologic unit 3 (fig. 24). Upward and downward movement of water in hydrogeologic unit 3 reflects changes in head with the changing thickness and transmissivity of this hydrogeologic unit. Vertical flow is downward in areas where the aquifer thickens and transmissivity increases because head decreases with depth; flow is upward in areas where the aquifer thins and transmissivity decreases because head increases with depth (fig. 24). Hydrogeologic unit 2 generally acts as an impediment to the downward movement of water from the fractured basalt of hydrogeologic unit 1 because of the smaller hydraulic conductivity of hydrogeologic unit 2. The stratigraphic interpretation of Anderson and Liszewski (1997, fig. 27) indicates that southwest of the INL hydrogeologic units 1 and 2 were uplifted. As a result of uplift, hydrogeologic unit 1 is no longer saturated and hydrogeologic unit 2 intersects the water table (fig. 24). Hydrogeologic unit 2 is less transmissive than hydrogeologic unit 1 and thus the water-table gradient steepens from about 2 to 3 ft/mi (fig. 9) upgradient of where hydrogeologic unit 2 intersects the water table to about 25 to 30 ft/mi (Lindholm and others, 1988) at and downgradient of this location. The steepening of the gradient also may be caused by the addition of underflow to the aquifer from the Big Lost River, which significantly increases the volume of water moving through this area. The steeper gradient, increased aquifer transmissivity, and increased volume of water in this area probably facilitate the downward flow of water across the younger rocks of hydrogeologic unit 2 into the intermediate-age rocks of hydrogeologic unit 3.
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Department of the Interior | U.S. Geological
Survey
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