A helium isotope perspective on the Dixie Valley, Nevada, hydrothermal system

Geothermics 35 (2006) 26–43

A helium isotope perspective on the Dixie Valley, Nevada, hydrothermal system B. Mack Kennedy ∗ , Matthijs C. van Soest Center for Isotope Geochemistry, Earth Sciences Division, Lawrence Berkeley National Laboratory, MS 70A-4418, Berkeley, CA 94720-8179, USA Received 22 March 2005; accepted 8 September 2005 Available online 19 October 2005

Abstract Fluids from springs, fumaroles, and wells throughout Dixie Valley, NV were analyzed for noble gas abundances and isotopic compositions. The helium isotopic compositions of fluids produced from the Dixie Valley geothermal field range from 0.70 to 0.76 Ra, are among the highest values in the valley, and indicate that ∼7.5% of the total helium is derived from the mantle. A lack of recent volcanics or other potential sources requires flow of mantle-derived helium up along the valley bounding Stillwater Range Front Fault, from which the geothermal fluids are produced. Using a one-dimensional flow model, a lower limit fluid flow rate up through the fault of 7 mm/yr is estimated, corresponding to a mantle 3 He flux of ∼104 atoms m−2 s−1 . A comparison between the fluids from Dixie Valley springs, fumaroles, and wells and the fluids produced from the geothermal field reveals a mixing trend between the geothermal fluid and younger, cooler groundwaters. The exceptions are those features that either emanate directly from the Stillwater fault or wells that penetrate and extract fluids from the fault zone, all of which have helium isotopic compositions that are indistinguishable from the geothermal production fluids. The results of our study indicate that the Stillwater Range Front Fault system must act as a permeable conduit that can sustain high vertical fluid flow rates from deep within the crust and crust-mantle boundary and that high permeability may exist along most of its length. This suggests that the geothermal potential of the Stillwater fault may be significantly greater than the 6–8 km long system presently under production. Since all the numerous springs, wells, and fumaroles in the valley also contain a fluid component that is indistinguishable from the geothermal/Stillwater fault fluid, the potential for an additional deeper and more pervasive geothermal system also exists and should be further evaluated. Furthermore, we suggest that elevated helium isotope compositions in regions with little or no recent magmatism are an indicator of the deep crustal permeability that is required to drive and sustain extensional geothermal systems. © 2005 CNR. Published by Elsevier Ltd. All rights reserved. Keywords: Helium isotopes; Noble gases; Magmatic; Permeability; Geothermal; Dixie Valley; Basin and Range ∗

Corresponding author. Tel.: +1 510 486 6451; fax: +1 510 486 5496. E-mail address: [email protected] (B.M. Kennedy).

0375-6505/$30.00 © 2005 CNR. Published by Elsevier Ltd. All rights reserved. doi:10.1016/j.geothermics.2005.09.004

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1. Introduction The Basin and Range Province in western North America is characterized by an anomalous thermal gradient, high heat flux, volcanism, and extensional tectonics that have varied in time and space over the past ∼30 million years. This geologic environment has created a vast region of exceptionally high potential for geothermal energy development. Beginning ∼8 million years ago and continuing to the present, magmatism has migrated to the margins of the basin and NW–SE extension, characterized by high-angle block faulting, dominates in the northwestern and central Basin and Range (e.g. Parsons, 1995). Within the Province, two types of geothermal resource have been identified: magma-hosted and “extensional”. Concentrated along the margins of the Basin and Range, the magma-hosted systems (e.g. Steamboat Springs, Long Valley, Coso, and Roosevelt Hot Springs) mine heat from cooling magma bodies emplaced in the shallow crust and are similar to the numerous magma-hosted systems throughout the world. However, the extensional geothermal systems, which are found throughout the Basin and Range and may constitute the greatest geothermal potential, are nearly unique to this region. At present, very little is known, except in general terms, about the origin and development of the extensional systems or what drives their spatial and temporal distribution with respect to the regional and local tectonic history. The Dixie Valley geothermal field, located in the west central portion of the northern Basin and Range within a “sub-province” referred to as the central Nevada seismic zone, is widely considered to be a “classic” example of an extensional Basin and Range geothermal system (e.g. Sass, 1995). There is no evidence for an active or recently active magma system associated with Dixie Valley that is young enough to supply the requisite heat, but the valley and surrounding region is tectonically active and characterized by NW–SE extension (∼1 mm/yr over the last 12,000 years; Thompson and Burke, 1973; Caskey et al., 2000) and exceptionally high heat flow (∼80–90 mW/m2 ; Sass et al., 1994). The temperature of the production reservoir is ∼225–245 ◦ C (Benoit, 1992) and production depth now varies from ∼2.4 to 2.7 km. Primary fluid production is from large extensional fractures, associated with the Stillwater Range Front Fault system, which is a high-angle block fault system that is characteristic of the extensional style of faulting that has dominated the Basin and Range since mid to late Miocene. Local thermal gradients indicate that the Range Front Fault system also serves as the primary conduit for heat transport to the surface (Blackwell et al., 2000, 2002). Late Miocene basalts (K-Ar age of ∼8.5 Ma; Waibel, 1987) represent the most recent volcanic activity in the Dixie Valley area and, as such, are too old to provide an adequate heat source for the geothermal field. The lack of any geologic or geophysical evidence for an active or recently active nearsurface magma chamber implies that deep fluid circulation must be the driving force for the Dixie Valley geothermal system (e.g. Sass, 1995). Surprisingly, despite the regional nature of the high thermal gradient and the occurrence of warm springs and paleo-sinter and travertine deposits throughout Dixie Valley, the geothermal reservoir currently under production remains restricted to a narrow 6–8 km long zone adjacent to the range front fault. However, it is generally believed that the resource may be much more extensive (David Blackwell, 2005, pers. communication). In conjunction with a regional survey of the Dixie Valley hydrologic system (Goff et al., 2002), we collected water and gas samples from accessible wells within and outside the productive geothermal field, and from springs and fumaroles throughout Dixie Valley. The samples were analyzed for noble gas abundances and isotopic compositions to (1) determine fluid and heat

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source, (2) identify fluid flow paths and (3) evaluate the possibility for a more extensive deep geothermal reservoir. 2. Experimental procedures and results Samples of non-condensable gas or water were collected from wells, springs and fumaroles throughout Dixie Valley. Three types of wells were sampled: (1) high-temperature geothermal production wells in which boiling and phase separation occurs within the well bore; (2) lowto moderate-temperature wells in which a separate gas phase forms upon decompression as the fluid flows up the well bore, and (3) low- to moderate-temperature wells that produce a single liquid phase. In sampling Type 1 wells, the wellhead vapor phase was passed sequentially through water and ice-bath condensers. The cooled fluid flows through an inverted “Y-shaped” tube separating the condensate and gas phases, allowing for collection of the non-condensable gases at ambient conditions. When sampling Type 2 wells, the water and ice-bath condensers were bypassed and the two-phase wellhead fluid was passed directly through the inverted “Y-shaped” separator, isolating a sample of the free-gas phase for collection. In both cases, the entire sampling apparatus was initially filled with either the wellhead liquid or the liquid condensate; the separated gas phase was captured using a water displacement technique. The single-phase liquid flowing from Type 3 wells was sampled directly for analysis of the dissolved gases. Springs that were either non-boiling or without a free-gas phase were sampled in much the same way as the Type 3 wells. A tube was submerged in the spring, as close to the spring inlet as possible, and water was pulled through the sampling apparatus using a peristaltic hand pump. After several volumes of water had passed through the system a sample of the liquid was collected for analysis of the dissolved gases. When sampling boiling springs or springs with a free-gas phase, the vapor/gas phase was captured and forced through the water-filled sampling apparatus using an inverted funnel submerged in the spring. The non-condensable gas phase was separated and collected from the liquid phase using the “Y-shaped” separator. Water displacement was used to capture the gas phase in the sample container, in the same manner as used for well Types 1 and 2. Fumarole gases were captured and forced through the sampling apparatus using an inverted funnel sealed around the funnel-ground contact. The gas and vapor are allowed to flow through the sampling apparatus in order to adequately purge the sampling lines of air before collecting a sample. In all cases, a 9.8 cm3 sample of gas or liquid was collected at ambient conditions in a Cu-tube cold-welded at each end using bolt-driven clamps. To ensure sample integrity, the clamps remained in place until the sample was ready for analysis and attached to the sample preparation vacuum line, which is in series with the noble gas mass spectrometer. Sample preparation and the noble gas analyses were conducted in the Roving Automated Rare Gas Analysis (RARGA) laboratory at the Lawrence Berkeley National Laboratory. Sample preparation, analytical techniques and the instrumentation used in the RARGA laboratory are similar to those described by Kennedy et al. (1985), Smith and Kennedy (1985) and Hiyagon and Kennedy (1992). The primary difference between this and earlier studies is in the sample preparation technique. In our earlier version of the sample preparation line, CO2 was removed from the gas stream by exposure to 4N NaOH, pre-degassed and under vacuum, and the water vapor was removed by exposure to CaO powder at room temperature. In the newest version of our sample preparation line, these steps have been replaced with more time-efficient procedures. Water vapor is condensed in a flow through a trap cooled externally using a methanol-liquid N2 slurry. CO2 and other reactive gases

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(N2 , H2 , CO, etc.) are chemically removed by exposure to a stream of evaporating Ti-metal. After removal of the reactive gases, an aliquot of the remaining purified noble gas fraction is isolated for determination of absolute and relative abundances. The rest of the noble gas fraction (∼95%) is trapped on activated coconut charcoal cooled to ∼35 K, from which each noble gas can be thermally separated and analyzed individually for its isotopic composition. The sample preparation line and mass spectrometer performance are calibrated using aliquots of air and a Berkeley helium standard with an isotopic composition of 2.4 Ra (where Ra is the 3 He/4 He ratio in air: 1.4 × 10−6 ). Noble gas relative abundances are presented in Table 1. The relative abundances are given in F-value notation in which measured relative abundances are normalized to the air abundance with 36 Ar as the reference isotope [e.g. F(i) = (i/36 Ar)sample /(i/36 Ar)air ]. As defined, F-values are fractionation factors that provide a measure of enrichment or depletion relative to atmospheric composition. The relative abundance data for the high-temperature production wells (Type 1, see above) have been reconstituted to reservoir conditions. In all cases, the applied corrections were minor, owing to the very low solubility of noble gases in water at the wellhead temperatures. Any potential uncertainty in the corrected abundances, reflecting uncertainties in steam fractions, gas/steam ratios, enthalpy, etc. will have little or no bearing on the subsequent discussions. For easy comparison, the F-value compositions of air noble gases dissolved in water at various temperatures are also tabulated. The helium R/Ra values, where R is the measured sample 3 He/4 He ratio and Ra is the ratio in air (1.4 × 10−6 ), are also summarized in Table 1. Argon isotopic compositions were determined, but in all cases the compositions were indistinguishable from air and therefore they are not included in the data table. This paper will focus primarily on the helium relative abundances [F(4 He)] and isotopic compositions (R/Ra). A detailed discussion of the F-values [other than F(4 He)] as related to the Dixie Valley reservoir, such as their importance in monitoring re-injected production fluids, can be found in Kennedy et al. (1999) and Kennedy and Shuster (2000). 3. Discussion The noble gases (He, Ne, Ar, Kr, and Xe) and their isotopes are excellent natural tracers for fluid source and migration in the Earth’s crust (e.g. Lupton, 1983; Kennedy et al., 1985; Hiyagon and Kennedy, 1992; Kennedy and Truesdell, 1996). Due to their inert chemical character, they provide long-lasting tracers that can see through complex chemical processes affecting the compositions of other more reactive species. A geothermal reservoir fluid can contain noble gases from mantle, crustal and atmospheric sources. As each of these sources is characterized by a unique noble gas composition, their contributions to a geothermal reservoir can be readily identified and used to constrain a fluid’s history. Helium isotopes are of particular interest as they provide unequivocal evidence for the presence of mantle-derived volatiles in geothermal systems, and are therefore an indication of heat source and of the role that mantle melting plays in the formation of a crustal geothermal system. Helium associated with crustal fluids that have experienced no mantle influence is dominated by radiogenic 4 He produced from radioactive decay of U and Th to Pb and is characterized by a 3 He/4 He ratio of ∼0.02 Ra. Helium associated with mantle fluids is enriched in 3 He acquired during earth formation. Helium in mid-ocean ridge basalts worldwide is characterized by a 3 He/4 He ratio of ∼8–9 Ra, which is generally believed to represent the upper mantle composition. Ratios as high as ∼35 Ra have been observed in ocean island basalts thought to be related to deep “plume” volcanism. Subduction zone volcanism is characterized by ratios of ∼6–9 Ra. Mantle helium

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Table 1 Noble gas relative abundances and helium isotopic composition of fluids from Dixie Valley, NV Sample ID

Separator/well

Map #

Latitude (N)

Longitude F(4 He) (E)

±

F(22 Ne)

±

F(84 Kr)

±

F(132 Xe)

±

R/Ra

±

Geothermal field Sample year: 1995 DV95-01 V102 + 103 DV95-02 76A-7(V104) DV95-03 V105 DV95-04 V101

1 2 3 4

I I I I

39.9639 39.9590 39.9663 39.9869

242.1418 242.1430 242.1463 242.1693

159.06 150.18 122.93 149.86

9.69 9.11 7.46 9.09

0.3763 0.4464 0.4916 0.3985

0.0134 0.0115 0.0131 0.0110

1.8007 1.6401 1.6403 1.6724

0.0442 0.0324 0.0364 0.0331

3.8710 2.8350 2.6530 2.7110

0.2747 0.1563 0.3460 0.1524

0.76 0.70 0.70 0.72

0.03 0.03 0.05 0.05

Sample year: 1997 DV97-1G 73-7 DV97-3G 76A-7(V104) DV97-10G 45-33 DV97-4G 73B-7 DV97-6G 82A-7 DV97-5G 74-7 DV97-7G 28-33 DV97-2 63-7 DV97-9SG V101

5 2 6 7 8 9 10 11 4

I I I I I I I I I

39.9655 39.9590 39.9897 39.9652 39.9675 39.9637 39.9852 39.9650 39.9869

242.1424 242.1430 242.1740 242.1445 242.1452 242.1420 242.1682 242.1420 242.1693

142.41 156.13 136.30 103.70 140.63 147.07 134.25 150.55 135.14

3.21 3.56 3.08 2.34 3.16 3.32 3.02 3.46 3.05

0.5397 0.3905 0.4182 0.5596 0.4576 0.4547 0.3883 0.4239 0.3955

0.0324 0.0233 0.0255 0.0332 0.0273 0.0271 0.0229 0.0256 0.0237

1.6898 1.7779 1.7784 1.5792 1.7401 1.7821 1.7736 1.7747 1.8115

0.0180 0.0172 0.0238 0.0199 0.0177 0.0178 0.0284 0.0176 0.0197

2.8347 3.1653 3.2651 2.5577 3.0207 3.0340 3.0874 3.0099 3.2828

0.0789 0.0709 0.0635 0.0671 0.0669 0.0712 0.0677 0.0550 0.0851

0.77 0.66 0.72 0.65 0.70 0.69 0.71 0.63 0.68

0.03 0.03 0.04 0.03 0.03 0.03 0.03 0.03 0.03

Sample year: 1998 DV98-02 27–33 DV98-03 37–33 DV98-04 28–33 DV98-05 76A-7(V104) DV98-06 74-7 DV98-07 63-7 DV98-08 73-7 DV98-09 82A-7 DV98-11 73B-7

12 13 10 2 9 11 5 8 7

I I I I I I I I I

39.9869 39.9861 39.9852 39.9590 39.9637 39.9650 39.9655 39.9675 39.9652

242.1693 242.1688 242.1682 242.1430 242.1420 242.1420 242.1424 242.1452 242.1445

135.66 135.23 188.67 198.00 239.09 152.80 10.40 175.00 179.00

3.08 3.04 4.27 9.88 6.09 3.45 0.52 8.77 8.95

0.2831 0.3205 0.3064 0.3880 0.3662 0.3427 0.9300 0.3620 0.3780

0.0392 0.0452 0.0563 0.1000 0.0627 0.0463 0.0253 0.0682 0.0928

1.8734 1.7652 1.9473 1.7500 1.5665 1.7618 1.0400 1.7600 1.8000

0.0201 0.0178 0.0198 0.0354 0.0186 0.0209 0.0212 0.0379 0.0394

3.4781 3.0486 3.9892 2.9900 2.4452 3.0243 1.0600 3.0600 3.0800

0.0633 0.0626 0.0758 0.1560 0.0598 0.0601 0.0559 0.1550 0.1590

0.73 0.73 0.68 0.78 n.d. 0.71 0.69 0.68 0.67

0.04 0.04 0.03 0.03 n.d. 0.04 0.03 0.03 0.04

B.M. Kennedy, M.C. van Soest / Geothermics 35 (2006) 26–43

Type

Well # Goeringer Goeringer 27-32 46-32 45-14 66-21 36-14

14 14 15 16 17 18 19

II, III II, III II, III II, III II, III II, III II, III

39.9693 39.9693 39.9862 39.9881 39.8659 39.9311 39.9474

242.1409 242.1409 242.1516 242.1566 241.9951 242.0716 242.0926

9.64 8.72 121.00 130.00 86.85 100.00 148.64

0.48 0.11 6.05 6.53 2.17 5.01 15.55

0.3360 0.3258 0.3600 0.2960 0.2719 0.0669 0.3367

0.0650 0.0048 0.0556 0.0358 0.0279 0.1790 0.0206

1.8200 1.7640 1.9200 2.0600 1.8840 2.6200 1.8301

0.0372 0.0087 0.0481 0.0420 0.0194 0.0567 0.0167

3.2100 3.0281 3.6500 4.2200 3.5048 6.6800 3.0435

0.1620 0.0410 0.1910 0.2160 0.0747 0.3470 0.1562

0.40 n.d. 0.58 0.58 0.59 0.31 0.77

0.05 n.d. 0.04 0.02 0.03 0.06 0.03

Springs DV97-11 DV97-12 DV97-13 DV97-14 DV98-19 NV03DXHS NV03MCHS NV03MCHW NV03JHS1 NV03JHS2

Spring name Senator Senator’s Toe Sou Spring (G) Hyder Spring (G) Section 10 Fumarole Dixie Meadows (W) McCoy Hot Spring (G) McCoy Hot Spring (W) Upper Jersey Seep (W) Jersey Hot Spring (G)

20 21 22 23 24 25 26 26 27 28

V V IV IV V IV IV IV IV IV

39.9945 39.9872 40.0890 40.0035 39.9541 39.8004 40.0795 40.0795 40.1779 40.1782

242.1480 242.1568 242.2760 242.2831 242.0828 241.9408 242.3964 242.3964 242.5109 242.5042

70.24 45.52 60.67 32.64 5.09 13.24 6.85 3.25 28.13 32.77

1.58 1.03 1.37 0.73 0.26 0.15 0.08 0.04 0.33 0.38

0.4264 0.2808 0.4401 0.3964 1.0200 0.4028 0.7697 0.3758 0.3958 0.6212

0.0509 0.0686 0.0269 0.0291 0.0218 0.0075 0.0065 0.0047 0.0080 0.0098

2.0200 2.1075 1.5373 1.6454 0.9980 1.9094 1.1595 1.7735 1.8128 1.3407

0.0234 0.0228 0.0153 0.0263 0.0200 0.0155 0.0085 0.0134 0.0136 0.0096

5.8321 5.7278 2.0671 2.9683 1.0400 4.0710 1.5089 3.5601 3.5896 1.8092

0.1088 0.1057 0.0360 0.0667 0.0530 0.2271 0.0878 0.1980 0.2009 0.1031

0.66 0.69 0.53 0.44 0.86 0.84 0.30 0.34 0.52 0.47

0.06 0.04 0.02 0.02 0.04 0.04 0.03 0.04 0.02 0.02

F(4 He) Reference values for air-saturated water (ASW) and air 10 ASW 0.22 20 ASW 0.26 Air 1.00

F(22 Ne)

F(84 Kr)

F(132 Xe)

R/Ra

0.2722 0.3090 1.0000

1.9412 1.8402 1.0000

3.677 3.279 1.000

1.000 1.000 1.000

Type I: high-temperature geothermal production well with phase separation; Type II: low-to-moderate temperature well with a separate gas phase upon decompression; Type III: low-to-moderate temperature well with a single liquid phase; Type IV: hot spring with (G) or without (W) a free gas phase; Type V: fumarole.

B.M. Kennedy, M.C. van Soest / Geothermics 35 (2006) 26–43

Valley wells DV98-12 DV99-01 DV98-15 DV98-16 DV98-17 DV98-18 DV02-01

31

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Fig. 1. Range in helium isotopic compositions (given as R/Ra) found in a variety of crustal hydrothermal systems. At one extreme is the upper mantle as defined by Mid-Ocean Ridge Basalts (MORB). At the other is helium produced in the crust from radioactive decay of U and Th. Intermediate values represent either mixtures of the two end-members or aging (see text).

enters the crustal hydrologic system either by direct intrusion and degassing of mantle-derived magmas or by the invasion of geo-pressured fluids degassed from deep mantle melts. After mantle fluids are injected into the crust, the original mantle helium isotopic composition will become diluted by the addition of crustal radiogenic 4 He, lowering the 3 He/4 He ratio. The degree to which the ratio is lowered (diluted) will depend on the flux of mantle helium into the crust, the production rate of radiogenic 4 He in the crust, which is proportional to the concentrations of U and Th, and the residence age or fluid flow rate within the crust. For instance, the helium isotopic composition in crustal fluids that mine heat from active near-surface magmatic systems (Fig. 1), typically, are similar to the composition in the mantle source [e.g. The Geysers (6–9 Ra); Kennedy and Truesdell, 1996]. In contrast, smaller and more evolved systems show clear evidence for dilution by crustal helium [e.g. Roosevelt Hot Springs (∼2 Ra); Welhan et al., 1988]. 3.1. Evidence for a mantle component at Dixie Valley The helium associated with the production fluids from the Dixie Valley geothermal reservoir has an isotopic composition of 0.70–0.76 Ra (Table 1), which are among the highest ratios measured throughout the valley. The ratios are elevated with respect to those expected for a purely crustal source (R ∼ 0.02 Ra), indicating unequivocally that a magmatic or mantle helium component is present in the Dixie Valley reservoir fluid. If it is assumed that the mantle source has a helium isotopic composition of 8–9 Ra, then as much as ∼7.5% of the total reservoir helium is mantlederived (Kennedy et al., 2000).

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The mantle signature, however, is not very strong when compared to geothermal systems that are known to be associated with recent igneous activity (Fig. 1). For instance, at Steamboat Springs, Nevada, 3 He/4 He ratios of ∼6 Ra (Torgersen and Jenkins, 1982; Welhan et al., 1988) have been measured, providing strong evidence for an active or recently active shallow magma system. Within the Basin and Range, high ratios reflecting active magmatic systems have also been reported for Long Valley and the Coso geothermal systems (Sorey et al., 1993; Welhan et al., 1988). The much lower Dixie Valley ratios, however, imply a lower 3 He flux and are inconsistent with a young shallow crustal magma heat source, confirming the lack of geologic and geophysical evidence for a near-surface magma system. Never the less, mantle 3 He is present in the Dixie Valley fluids. Given the absence of an active degassing crustal magma system, the potential 3 He sources are limited to the following: (1) fluid circulation through either the erupted Miocene basalts or an aged and non-active magma chamber originally charged with mantle helium, i.e. the source chamber for the Miocene basalts; or (2) fluid transport up along the range-front fault from much deeper sources, either a deep melt zone at or near the base of the crust (e.g. Jarchow et al., 1993) or fluids exsolved from deeper melting (approximately hundreds of km) that have found a pathway into the range front fault system (e.g. Kennedy et al., 1997). In the first case, the initial 3 He/4 He ratios in the basalt and/or magma chamber would be high and similar to the mantle source. Once the chamber, or erupted basalt, is isolated and its volatile inventory is no longer being replenished by the mantle source, the 3 He/4 He ratio will decline with time due to the addition of radiogenic 4 He from U and Th radioactive decay, resulting in much lower ratios today. In the second case, the high 3 He/4 He ratios associated with the mantle source would be diluted by the addition of radiogenic helium derived from surrounding crustal rocks, added to the fluid either by diffusion or mineral dissolution while the fluid is in transit through the range front fault system from the deep source to the reservoir. Both potential 3 He sources are evaluated below. 3.1.1. Fluid circulation through erupted or intruded Miocene basalts As fluids circulate through and equilibrate with the crust, the isotopic compositions of dissolved constituents will be modified by water-rock processes such as mineral dissolution, diffusive exchange, etc. Geothermal water derived from deep circulating groundwater will acquire an isotopic composition that reflects the water-rock processes and rock compositions integrated over the flow path of the fluid. In the absence of fluid sources other than deep circulating groundwater, one potential source for the excess (mantle-derived) 3 He in the Dixie Valley thermal waters is circulation and isotopic equilibration with the Miocene basalts or their intruded equivalents. Although the helium isotopic composition of the Dixie Valley basalts has not been determined, a comparison with other Basin and Range basaltic volcanic systems, for which there are He data (Reid and Graham, 1996; Dodson et al., 1998), provides a strong argument that over an extended period of time (several million years), erupted and emplaced basalts cannot retain high R/Ra values due to their very low helium concentrations and post-eruption/emplacement production of radiogenic 4 He. The time-dependent change in the helium isotopic composition of a reservoir, due to the addition of radiogenic 4 He, is given by
3

He 4 He

 = t


3

He 4 He

  0

1 1 + [4 He∗ ]t /[4 He]0



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where subscripts ‘t’ and ‘o’ refer to present-day and initial concentrations or ratios, respectively, and the time-dependent concentration of radiogenic 4 He ([4 He* ]t ) is given by     235 U 4 ∗ 238 [ He ]t = [ U]t [8 exp(λ238 t) − 1] + 7 238 exp(λ235 t) − 1 U 
232  

Th + 6 238 exp(λ232 t) − 1 U For times much less than the half-lives of U and Th (t1/2 ∼ 1010 years), this equation can be reduced to a simpler form given by
232 

 Th 4 ∗ 238 −7 −8 [ He ]t = [ U]t t 1.207 × 10 + 2.797 × 10 238 U The 238 U concentration is in grams/gram and the 4 He concentrations are in cm3 STP/g [STP: standard temperature (0 ◦ C) and pressure (1 bar)]. The ability of the emplaced or erupted basalt to retain a helium isotopic composition similar to its initial value for an extended period of time depends strongly on the post-eruption/emplacement (238 U/4 He) ratio. This ratio has not been measured in the Dixie Valley basalts. However, based on a variety of observations made on other systems, it is expected that this ratio is very high. For instance, it has been estimated that more than 90% of dissolved CO2 in basaltic melts is lost during pre-eruptive degassing, with an additional ∼6% lost during eruption (Zindler and Hart, 1986). Since He distribution coefficients between basaltic melt and vapor are significantly lower than that for CO2 (e.g. Jambon et al., 1986), emplacement and eruption will result in a dramatic decrease in dissolved 4 He and concomitant increase in the (238 U/4 He) ratio. Since (238 U/4 He) enrichment factors of ∼106 or more can be expected, this would lead to a rapid decline in the postemplacement/eruptive helium isotopic composition. Fig. 2 provides three illustrative examples. In each, the following assumptions have been applied: (3 He/4 He)0 = 8 Ra, (238 U) = 1 ppm, and (232 Th/238 U) = 3. The initial helium isotopic composition is consistent with that in sub-continental mantle, as estimated from analyses of Basin and Range basalts (Reid and Graham, 1996; Dodson et al., 1998). The U and Th concentrations are comparable to those of the average crust. The three examples differ only in the assumed initial (post-emplacement/eruption) 4 He concentrations of 10−9 , 10−8 , and 10−7 cm3 STP/g. The selection of an assumed initial 4 He concentration is somewhat arbitrary. The lowest concentration (10−9 cm3 STP/g) was chosen to be consistent with concentrations of mantle-derived helium measured in erupted basalts throughout the Basin and Range Province (Reid and Graham, 1996; Dodson et al., 1998) and is likely to be representative of the Miocene basalt occurring throughout Dixie Valley. The highest initial 4 He concentration (10−7 cm3 STP/g) was chosen as a reasonable upper limit for non-degassed basalt melts determined from mafic xenoliths and nodules entrained in continental volcanics (Mamyrin and Tolstikhin, 1984). However, it is not clear that 4 He concentrations in these xenoliths are representative of the sub-continental basaltic melts. It has been argued that the high concentrations reflect contamination by the invasion of fluids exsolved from partial melts deeper in the mantle (e.g. Dunai and Porcelli, 2002); these are included here as an example representing a strict upper limit for the 4 He concentration. Given the very low concentrations expected in erupted basalts (∼10−9 cm3 STP/g), it is reasonable to assume that the 8–9 My old Dixie Valley basalt occurring throughout Dixie Valley is not a viable source for the excess 3 He in the geothermal fluids. As demonstrated in Fig. 2, after only ∼1 My,

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Fig. 2. Change in helium isotopic composition (given as R/Ra) of a closed reservoir, such as an erupted basalt or isolated crustal magma chamber, due to the addition of radiogenic helium from the in situ decay of U and Th. Each curve 4 He concentrations in cm3 STP/g represents a different initial helium concentration.

the addition of radiogenic 4 He will reduce the helium isotopic composition of a typical erupted basalt to <0.1 Ra. High 3 He/4 He ratios consistent with the geothermal fluid would require fluid circulation through systems with 4 He concentrations greater than ∼10−7 cm3 STP/g, two orders of magnitude higher than have been observed in Basin and Range volcanics. Since ∼90% of the CO2 and >90% of the helium (due to the lower melt/vapor distribution coefficient of helium) in a partial melt are lost during pre-eruptive degassing, a comparable argument can be made for the source magma chamber or other intruded equivalents of similar age. Furthermore, as will be discussed below, there is evidence that the 3 He-carrying fluid component is not restricted to the producing geothermal reservoir, but instead occurs throughout Dixie Valley. It seems unlikely that sparsely distributed erupted basalt or an isolated magma chamber could influence such a vast hydrologic system. We therefore feel confident that neither the erupted Miocene basalts nor related intrusives can be the source of the mantle 3 He. 3.1.2. Upflow of deep mantle fluids Fluids circulating through young Quaternary extensional basins commonly contain 3 He in excess of that expected for radiogenic production, and in some cases the excess 3 He concentration is found to increase with depth in the basins (Ballentine et al., 1991, 2002). It is believed that the source of the excess 3 He is an enhanced flux of mantle volatiles up into the basin due to crustal thinning, crustal underplating by mantle melts, and/or similar processes related to extensional regimes (e.g. O’Nions and Oxburgh, 1983; Ballentine et al., 1991, 1996). Excess 3 He has also been found in fluids from seismically active areas undergoing compression. For example, fluids associated with the San Andreas Fault system in California, where the direction of principle stress is nearly perpendicular to the fault strike, have 3 He/4 He ratios as high as 4 Ra, despite a lack of any recent magmatism. It is believed that the excess 3 He is carried into the San Andreas

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Fig. 3. Simple model depicting focused flow of mantle volatiles through a ductile lower crust into a range-bounding normal fault, like the Stillwater Range Fault system, Dixie Valley, NV. Fluids leave the mantle with a helium isotopic composition of ∼8–9 Ra. During transit through the crust, the helium isotopic composition is diluted with crustal helium (0.02 Ra), creating a gradient in the fault zone helium isotopic composition that, at steady state, is proportional to the fluid flow rate through the fault.

system by geo-pressured mantle fluids focused through the ductile lower crust and into the roots of the fault system, and that the fluids may serve as a weakening agent for strike-slip fault rupture (Kennedy et al., 1997). In a similar manner, focusing of mantle volatiles into the shallow crustal fluid system may be at work in the extensional regime of Dixie Valley and associated Stillwater Range Front Fault system. A simple model for fluid flow up through the Range Front Fault is shown schematically in Fig. 3. Decompression during diapiric rise of mantle material will result in deep partial melting and may lead to crustal underplating. Geo-pressured mantle volatiles, such as CO2 and helium, released during partial melting, may be focused through weak zones in the ductile lower crust that are associated with the roots of the Range Front Fault. Once through the ductile region, the fluids move freely up along the fault system, invading the shallow crust and interacting with the basin hydrology. Initially, the helium released by partial melting and focused into the roots of the fault system will have an isotopic composition similar to that in the mantle source, presumably ∼8–9 Ra. [Even though the helium isotopic composition of subcontinental asthenosphere has not been measured directly, it is thought to be on the order of 8 Ra (Reid and Graham, 1996; Dodson et al., 1998)]. While flowing to the surface through the Range Front Fault system, the mantle helium isotopic composition will become diluted with radiogenic helium (0.02 Ra) derived from the ambient crustal rocks. The dilution of the mantle component with radiogenic helium will produce a depth-dependent gradient in the helium isotopic composition of the rising fluid. Using a one-dimensional steady-

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state advection model (no dispersion), the rate the helium isotopic composition decreases as the fluids rise to the surface will depend on the fluid flow rate through the fault zone, the scale length of transport, the rate radiogenic helium is added to the fluid, and the concentration of mantle helium in the fluid (e.g. Johnson and DePaolo, 1994; Kennedy et al., 1997). The steady-state upward flow rate (q), scaled to a characteristic length defined by the local crustal thickness (Hcrust ), is given by q=

  Hcrust ρs P(4 He) (R/Ra)meas − (R/Ra)crust ρf [4 He]f,mantle (R/Ra)mantle − (R/Ra)meas

where ρs and ρf refer to the density of the rock and fluid and (R/Ra) refers to the helium isotopic composition in the crust, mantle and measured in the fluid at the sampling point. Under steady-state conditions, the rate radiogenic helium is added to the fluid will be approximately equivalent to the present-day production rate [P(4 He)] and, therefore, dependent only on the average concentration of U and Th in the country rocks. The concentration of mantle helium in the fluid ([4 He]f,mantle ∼ 1.6 × 10−7 cm3 NTP/g fluid; ‘NTP’ refers to conditions at 1 atm and 20 ◦ C) can be calculated assuming that ∼7.5% of the measured reservoir fluid helium content (∼21 × 10−7 cm3 NTP/g fluid) is derived from the mantle. Using a scale length for transport equivalent to the local crustal thickness (∼30 km; Jarchow et al., 1993), gives a vertical fluid flow rate through the fault zone on the order of ∼7 mm/yr. This rate is a strict lower limit, since it is calculated assuming one-dimensional advective flow with no diffusive re-distribution or mixing with other fluids that may be enriched in radiogenic 4 He. The latter assumption is highly unlikely. Chemical and isotopic compositions of surface fluids throughout Dixie Valley are consistent with a derivation from local Pleistocene meteoric and lake waters that have experienced variable and extensive water-rock exchange (Goff et al., 2002; Nimz et al., 2004). These initially cool fluids circulated deep down into the fault system and were heated by the prevailing subsurface temperature, eventually rising back to the surface driven by thermal buoyancy. In transit these fluids will acquire radiogenic 4 He along their flow path and when mixed with the rising deeper mantle fluids their 3 He/4 He ratio will be further diluted. Obviously, the assumptions that go into the model calculation impart a fair degree of uncertainty. Never the less, the flow rate of deep fluids into the reservoir is likely to be significantly greater than ∼7 mm/yr due to probable mixing with other non-3 He bearing fluids but slower than flow within the reservoir (∼5–120 m/day, as estimated from tracer tests; Adams et al., 1989; Benoit, 1992). The much lower deep fluid flux into the reservoir, relative to flow within the reservoir, is consistent with the steady decline in the proportion of indigenous reservoir fluid relative to re-injected fluids produced from the geothermal field over time (Benoit, 1992; Kennedy et al., 1999). The lower limit fluid flow rate through the fault zone, of ∼7 mm/yr, is equivalent to a 3 He flux of ∼104 atoms m−2 s−1 (using the same parameters given above for solving the flow rate equation). Helium and heat in the Earth are thought to be uniquely coupled, with 4 He production from U and Th decay accounting for ∼70–75% of the terrestrial heat budget (e.g. O’Nions and Oxburgh, 1983; Kennedy et al., 2000). Using measured 3 He and seawater temperature anomalies in fluids related to mid-ocean ridge hydrothermal systems, a lower limit mantle 3 He–heat ratio of ∼3500 3 He atoms/mW s has been estimated (Jenkins et al., 1978; see also Torgersen, 1993 and references therein). Assuming this value to be representative of the sub-continental mantle, the 3 He flux implies a mantle heat flux of ∼3 mW m−2 , which is ∼8 times lower but comparable to the estimated residual (crustal contribution subtracted) heat flux of 20–30 mW m−2 that is typical of the Dixie Valley area (Sass et al., 1994). Although the many assumptions that go into

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Fig. 4. Shaded relief map of northern Dixie Valley and surrounding mountain ranges, showing numbers of sampling locations, which correspond to the numbers under the heading Map # in Table 1. GF = Dixie Valley geothermal field. Black triangles: wells (Types I–III from Table 1) sampled for this study; black circles: springs (Type IV) sampled for this study; white stars: fumaroles (Type V) sampled for this study. White and black dotted circles and white asterisks indicate other features (springs, wells, and streams) that were sampled by Goff et al. (2002). The northwest corner of the map has a latitude/longitude location of 40.1705◦ N and 241.8970◦ E, and the southeast corner 39.7475◦ N and 242.5625◦ E.

this calculation introduce significant uncertainty, the relatively close agreement between the 3 Hecalculated heat flux with that determined directly from measured temperature gradients corrected for crustal production seems to imply that mantle melting contributes to the overall heat flux. Without partial melting there would be no 3 He flux and the heat would have to be released by some other mechanism, such as solid state underplating of mantle material. The presence of mantle helium in the Dixie Valley fluids requires the ascent of mantle volatiles through the range front fault system. The helium isotopic compositions observed in Dixie Valley fluids depend on the fluid flow rate through the fault system, the extent of mixing with downward percolating surface-derived fluids, the local crustal thickness, which sets limits on the length of the fluid flow path, and the U and Th concentrations integrated over the flow path. Assuming the latter two variables are relatively constant within the valley, the helium isotopic compositions of the Dixie Valley fluids require a highly permeable Stillwater fault system that can sustain high vertical fluid flow rates to the surface from deep in the crust and the crust mantle boundary. As a potential exploration tool, helium isotopes may provide the only indicator of the deep crustal permeability needed to drive and sustain convective fault-hosted hydrothermal systems. 3.2. Helium evidence for a more extensive geothermal system The Dixie Valley geothermal reservoir fluids contain excess 3 He derived from the mantle and delivered to the geothermal reservoir through deep permeable pathways provided by the Stillwater Range Front fault system. The valley is approximately 120 km long and 20–30 km wide, defined on the west by the Stillwater Range and the east by the Clan Alpine Range. Throughout the valley and within the bordering ranges there are numerous springs, wells, and fumaroles (Fig. 4) which

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Fig. 5. Helium isotopic compositions of a variety of Dixie Valley fluids plotted as a function of the helium enrichment factor [F(4 He)]. Solid gray circles are hot springs and fumaroles, solid black circles are wells outside of the geothermal field. The ellipse labeled DVGF represents the range in compositions found in the production fluids from the Dixie Valley geothermal field (squares, triangles, and diamonds; symbols delineate year of sampling). The dashed line depicts mixing between the DVGF fluid and a younger, cooler shallow groundwater. The dotted line depicts boiling and/or air contamination (see text). The numbers in parentheses are upper limit estimates of water-rock equilibration temperatures (Goff et al., 2002).

were sampled and analyzed as part of the regional survey of the Dixie Valley hydrologic system (Goff et al., 2002). Many of these features have also been sampled and analyzed for noble gas isotopes and abundances, in part to provide a better overall understanding of the Dixie Valley hydrologic system but also to evaluate the potential for a more extensive geothermal system. The helium enrichment factors [F(4 He)] and helium isotopic compositions for all of the Dixie Valley samples analyzed are summarized in Fig. 5 (sample locations can be found in Fig. 4). The compositions of the Dixie Valley geothermal production wells, which plot within the ellipse in Fig. 5, have the highest helium enrichment factors (∼200) and are the most enriched in 3 He (R ∼ 0.8 Ra). As depicted by the dashed line, the helium compositions of the remaining springs and wells throughout the valley are mixtures of this enriched (deep) fluid and a less 3 He- and 4 He-enriched end-member fluid [R < 0.35 Ra, F(4 He) < 5]. The very low 4 He-enrichment factor and more radiogenic helium isotopic composition in the latter fluid suggest a young groundwater. Chemical geothermometers, although not very reliable for many of the sampled features (Fraser Goff, 2002, pers. communication), in general indicate an increase in water-rock equilibration temperature (values in parentheses, Fig. 5) that crudely tracks the increase in the proportion of the geothermal end-member fluid [increasing F(4 He)], suggesting that the less enriched end-member is both young and cool and therefore is probably a young shallow groundwater. (The composition of Well 66-21, which falls off the dashed mixing line, is suspect due to sampling difficulties that may have led to phase separation and gas fractionation.) A very important ramification of the mixing line is that it requires that all of the Dixie Valley features analyzed for noble gases have been influenced by the same deep fluid that provides excess mantle 3 He to the geothermal production fluid and that this deep fluid occurs throughout the valley.

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The compositions of nearby fumaroles (Section 10 Fumarole, Senator Fumarole and Senator’s Toe) and the Dixie Meadows hot spring group are exceptions to the simple mixing trend; plotting both outside the ellipse defined by the production reservoir and off of the dashed mixing line. Senator Fumarole emanates from the Stillwater fault adjacent to the geothermal field. Senator’s Toe is located at the base of a fan adjacent to the field and is an outflow point for boiling fluid that may be supplying Senator Fumarole (Bergfeld et al., 2001). Section 10 Fumarole is a set of weak fumaroles located along the front of the Stillwater Range ∼5–6 km southwest of the geothermal field and nearly adjacent to well 36-14, a very high temperature (275 ◦ C at ∼3 km) well that penetrates the Stillwater Fault and has a helium enrichment factor and isotopic composition that is identical to the production reservoir. The Dixie Meadows hot spring group consists of several springs and seeps emanating from alluvium and fan deposits near the southern end of the Humboldt Salt Marsh, ∼30 km southwest of the geothermal field. When sampled in 2003, the highest temperature measured at the surface was 86.5 ◦ C at a vent closest to the range front fault. The springs emerge from either an outflow plume rising along the Stillwater fault zone or a buried rampart fault beneath Dixie Meadows (Fraser Goff, 2003, pers. communication). In Fig. 5, the dotted line connecting the compositions of these features with the ellipse defined by compositions of production wells in the geothermal field depicts the compositional trajectory expected for either boiling and/or air contamination. Boiling will cause noble gas fractionation driven by solubility differences, resulting in a preferential loss of 4 He and retention of 132 Xe with respect to 36 Ar, i.e. decreasing F(4 He) and increasing F(132 Xe), with no change in the 3 He/4 He ratio, as is the case for Senator Fumarole, Senator’s Toe and Dixie Meadows [Table 1; note that the F(132 Xe) values for these features are ∼2 times greater than that expected for air–saturated water]. Contamination with air will drive the F-values of all the noble gases and the R/Ra values towards a value of 1, which is the case for the Section 10 Fumarole (Table 1). These are the only surface features within Dixie Valley that have 3 He/4 He compositions indistinguishable from the geothermal production fluids and are the only features that appear to produce fluids directly from the Stillwater fault. This implies that the Stillwater fault, at least along the section from the geothermal field southwest to Dixie Meadows (a distance of ∼30 km), serves as a permeable path for vertical fluid flow that is relatively unencumbered by the shallow cooler groundwaters that are affecting fluid compositions further out in the valley. If so, the geothermal potential of the Stillwater fault system may be significantly greater than the 6–8 km long system presently under production. The other features in the valley, which plot along the mixing line in Fig. 5, although influenced by local hydrology, must also contain a fluid component indistinguishable from the geothermal/Stillwater fault fluid. Whether or not this also implies a deeper pervasive geothermal system throughout Dixie Valley needs to be further evaluated. 4. Summary Helium isotopic compositions (0.70–0.76 Ra) elevated relative to that expected for radiogenic production in the crust (0.02–0.10 Ra) indicate that as much as ∼7.5% of the helium in the fluids from the Dixie Valley geothermal reservoir is mantle-derived. Due to a lack of geologic and geophysical evidence for current or recent magmatic activity and the inability of older magma systems to sustain high enough 3 He/4 He ratios, the most viable source for the mantle 3 He is fluid transport up through the Stillwater Range Front fault system, a system that appears to be in direct communication with the mantle. Using a simple one-dimensional steady-state fluid flow model, the helium content and isotopic composition imply vertical fluid flow rates from the mantle of ∼7 mm/yr. This is a strict lower limit to the fluid flow rate: the one-dimensional

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model does not consider diffusive re-distribution of helium or mixing with water containing only a crustal helium component and therefore enriched in 4 He. The flow rate corresponds to a 3 He flux of ∼104 atoms m−2 s−1 , which is equivalent to a mantle heat flux of ∼3 mW m−2 [assuming a mid-ocean ridge mantle 3 He/heat ratio of ∼3500 3 He atoms/mW s (Jenkins et al., 1978)]. The magnitude of the 3 He-derived mantle heat flux is comparable to the reduced heat flux estimated from temperature gradients and an assumed distribution of U and Th in the crust. The similarity of these two estimates suggests that the anomalous Dixie Valley heat flux is partially driven by mantle melting. Unfortunately, there are significant uncertainties associated with the conversion of a 3 He mass flux to a flux of mantle-derived heat, so this conclusion requires further testing. The fluid samples from the Dixie Valley geothermal field have the highest helium isotopic compositions (R ∼ 0.8 Ra) and helium enrichment factors [F(4 He) ∼ 200] in Dixie Valley. The high-temperature fluids are produced from fractures along the valley bounding Stillwater Range Front fault system. The helium compositions of springs and wells throughout Dixie Valley that are not in direct communication with the Stillwater Range Front fault are a mixture of this deep fluid with a younger less 4 He- and 3 He-enriched groundwater [R < 0.35 Ra, F(4 He) < 5]. The exceptions to the simple mixing trend are wells, fumaroles and springs located along the range front fault that do not show any noble gas evidence of being influenced by the shallow groundwaters implied by the mixing line. The Stillwater Range Front Fault, at least along the section from the geothermal field southwest to Dixie Meadows, is a permeable path for vertical fluid flow that is unencumbered by the local hydrology that affects fluids further out in the valley. This would suggest that the geothermal potential of the Stillwater Range Front Fault system may be significantly greater than the 6–8 km long system presently under production. The other features in the valley, which plot along the mixing line in Fig. 5, although influenced by local hydrology, also contain a fluid component indistinguishable from the geothermal/Stillwater Range Front Fault fluid. Whether or not this implies an additional deeper pervasive geothermal system should be further evaluated. Furthermore, the connection between the occurrence of deep permeability and elevated helium isotope compositions suggests that the latter can be used as a new exploration tool to find fault systems characterized by the deep crustal permeability needed to drive and sustain fault-hosted geothermal systems lacking a magmatic heat source. Acknowledgements The authors would like to thank Stu Johnson and Dick Benoit for their assistance and support, particularly with field logistics and access to relevant field data. Thanks are also due to Dave Shuster for assistance with sample collection and analysis, and to Dave Hilton and Dave Blackwell for reviews. This work was supported by the U.S. Department of Energy, Office of Basic Energy Sciences and Office of Geothermal Technologies, under contract DE-AC03-76SF00098. References Adams, M.C., Benoit, W.R., Doughty, C., Bodvarsson, G.S., Moore, J.N., 1989. The Dixie Valley, Nevada tracer test. Geotherm. Resour. Council Trans. 13, 215–220. Ballentine, C.J., O’Nions, R.K., Oxburgh, E.R., Horvath, F., Deak, J., 1991. Rare gas constraints on hydrocarbon accumulation, crustal degassing and groundwater flow in the Pannonian Basin. Earth Planetary Sci. Lett. 105, 229–246. Ballentine, C.J., O’Nions, R.K., Coleman, M.L., 1996. A Magnus Opus: helium, neon and argon isotopes in a North Sea oil field. Geochim. Cosmochim. Acta 60, 831–849.

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Ballentine, C.J., Burgess, R., Marty, B., 2002. Tracing fluid origin, transport and interaction in the crust. In: Porcelli, D., Ballentine, C.J., Wieler, R. (Eds.), Noble Gases in Geochemistry and Cosmochemistry. Reviews in Mineralogy and Geochemistry, 47. Mineralogical Society of America, Washington, DC, pp. 539–614. Benoit, D., 1992. A case history of injection through 1991 at Dixie Valley, Nevada. Geotherm. Resour. Council Trans. 16, 611–620. Bergfeld, D., Goff, F.E., Janik, C.J., 2001. Elevated carbon dioxide flux at the Dixie Valley geothermal field, Nevada; relations between surface phenomena and the geothermal reservoir. Chem. Geol. 177, 43–66. Blackwell, D.D., Gollan, B., Benoit, D., 2000. Temperatures in the Dixie Valley, Nevada, geothermal system. Geotherm. Resour. Council Trans. 24, 223–228. Blackwell, D.D., Leidig, M., Smith, R.P., 2002. Regional geophysics of the Dixie Valley area: example of a large Basin and Range geothermal system. Geotherm. Resour. Council Trans. 26, 519–523. Caskey, S.J., Bell, J.W., Slemmons, D.B., Ramelli A.R., 2000. Historic faulting and paleoseismology of the central Nevada seismic belt. In: Lageson, D.R., Peters, S.G., Lahren, M.M. (Eds.), Great Basin and Sierra Nevada, Geological Society of America Field Guide 2. Geological Society of America, Boulder, Colorado, USA, pp. 23–44. Dodson, A., DePaolo, D.J., Kennedy, B.M., 1998. Helium isotopes in lithospheric mantle: evidence from Tertiary basalts of the western USA. Geochim. Cosmochim. Acta 62, 3775–3787. Dunai, T.J., Porcelli, D., 2002. The storage and transport of noble gases in the subcontinental mantle. In: Porcelli, D., Ballentine, C.J., Wieler, R. (Eds.), Noble Gases in Geochemistry and Cosmochemistry. Reviews in Mineralogy and Geochemistry, vol. 47. Mineralogical Society of America, Washington, DC, pp. 371–409. Goff, F., Bergfeld, D., Janik, C.J., Counce, D., Murrell, M., 2002. Geochemical data on waters, gases, scales and rocks from the Dixie Valley region, Nevada (1996–1999). LA Report LA-13792. Los Alamos National Laboratory, Los Alamos, NM. Hiyagon, H., Kennedy, B.M., 1992. Noble gases in CH4 -rich gas fields, Alberta, Canada. Geochim. Cosmochim. Acta 56, 1569–1589. Jambon, A., Weber, H., Braun, O., 1986. Solubility of He, Ne, Ar, Kr and Xe in a basalt melt in the range 1250–1600 ◦ C: geochemical implications. Geochim. Cosmochim. Acta 50, 401–408. Jarchow, C.M., Thompson, G.A., Catchings, R.D., Mooney, W.D., 1993. Seismic evidence for active magmatic underplating beneath the Basin and Range Province, Western United States. J. Geophys. Res. 98, 22095–22108. Jenkins, W.J., Edmond, J.M., Corliss, B.C., 1978. Excess 3 He and 4 He in the Galapagos submarine hydrothermal waters. Nature 272, 156–159. Johnson, T.M., DePaolo, D.J., 1994. Interpretation of isotopic data in groundwater-rock systems: model development and application to Sr isotope data from Yucca Mountain. Water Resour. Res. 30, 1571–1587. Kennedy, B.M., Lynch, M.A., Smith, S.P., Reynolds, J.H., 1985. Intensive sampling of noble gases in fluids at Yellowstone. I. Early overview of the data; regional patterns. Geochim. Cosmochim. Acta 49, 1251–1261. Kennedy, B.M., Kharaka, Y.K., Evans, W.C., Ellwood, A., DePaolo, D.J., Thordsen, J., Ambats, G., Mariner, R.H., 1997. Mantle fluids in the San Andreas fault system, California. Science 278, 1278–1281. Kennedy, B.M., Janik, K., Benoit, D., Shuster, D., 1999. Natural geochemical tracers for injectate fluids at Dixie Valley. In: Proceedings of the 24th Workshop on Geothermal Reservoir Engineering, Stanford University, Stanford, CA, pp. 108–115. Kennedy, B.M., Fischer, T.P., Shuster, D.L., 2000. Heat and helium in geothermal systems. In: Proceedings of the 25th Workshop on Geothermal Reservoir Engineering, Stanford Geothermal Program Report SGP-TR-165, pp. 167– 173. Kennedy, B.M., Truesdell, A.H., 1996. The Northwest Geysers high-temperature reservoir: evidence for active magmatic degassing and implications for the origin of The Geysers geothermal field. Geothermics 25, 365–389. Kennedy, B.M., Shuster, D.L., 2000. Noble gases: sensitive natural tracers for detection and monitoring injectate returns to geothermal reservoirs. Geotherm. Resour. Council Trans. 24, 247–252. Lupton, J.E., 1983. Terrestrial inert gases: isotope tracer studies and clues to primordial components in the mantle. Annu. Rev. Earth Planetary Sci. 11, 371–414. Mamyrin, B.A., Tolstikhin, I.N., 1984. Helium Isotopes in Nature. Elsevier, Amsterdam, 273 pp. Nimz, G., Janik, C., Goff, F., Dunlap, C., Johnson, S., 2004. Regional chemical and isotope hydrology of Dixie Valley, Nevada: influences of geothermal circulation and Pleistocene pluvial recharge. Geochim. Cosmochim. Acta, submitted for publication. O’Nions, R.K., Oxburgh, E.R., 1983. Heat and helium in the earth. Nature 306, 429–431. Parsons, T., 1995. The Basin and Range Province. In: Olsen, K.H. (Ed.), Continental Rifts: Evolution, Structure, Tectonics. Developments in Geotectonics 25, Publication #264 of the International Lithosphere Program. Elsevier, New York, NY, pp. 277–324.

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