Oxygen isotope systematics in an evolving geothermal system: Coso Hot Springs, California

    Oxygen isotope systematics in an evolving geothermal system: Coso Hot Springs, California Thomas M. Etzel, John R. Bowman, Joseph N. Moore, John W. Valley, Michael J. Spicuzza, Jesse M. McCulloch PII: DOI: Reference:

S0377-0273(16)30215-3 doi:10.1016/j.jvolgeores.2016.11.014 VOLGEO 5967

To appear in:

Journal of Volcanology and Geothermal Research

Received date: Revised date: Accepted date:

22 July 2016 22 November 2016 22 November 2016

Please cite this article as: Etzel, Thomas M., Bowman, John R., Moore, Joseph N., Valley, John W., Spicuzza, Michael J., McCulloch, Jesse M., Oxygen isotope systematics in an evolving geothermal system: Coso Hot Springs, California, Journal of Volcanology and Geothermal Research (2016), doi:10.1016/j.jvolgeores.2016.11.014

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ACCEPTED MANUSCRIPT Oxygen Isotope Systematics in an Evolving Geothermal System: Coso Hot Springs, California Thomas M. Etzel1,2, John R. Bowman1, Joseph N. Moore2, John W. Valley3, Michael J. Spicuzza3, Jesse M. McCulloch4 Dept. of Geology and Geophysics, University of Utah, Salt Lake City, UT 84112; 2Energy & Geoscience, Institute, University of Utah, Salt Lake City, UT 84108; 3Dept. Of Geoscience, University of Wisconsin, Madison, WI 54706; 4Terra-Gen Operating Company, Little Lake, CA, USA

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ABSTRACT

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Oxygen isotope and clay mineralogy studies have been made on whole rock samples and feldspar separates from three wells along the high temperature West Flank of the Coso

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geothermal system, California. The reservoir rocks have experienced variable 18O/16O depletion, with 18O values ranging from primary values of +7.5 ‰ down to -4.6 ‰. Spatial patterns of

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clay mineral distributions in the three wells are not closely correlated with the distributions expected from measured, pre-production temperature profiles, but do correlate with spatial

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patterns of 18O/16O depletion, indicating that the stability of clay minerals in the three wells is a function of fluid-rock interaction in addition to temperature. Detailed 18O measurements in the three wells identify a limited number of localized intervals of extensive 18O/16O depletion. These intervals document localized zones of higher permeability in the geothermal system that have experienced significant fluid infiltration, water-rock interaction and oxygen isotopic exchange with the geothermal fluids. The local zones of maximum 18O/16O depletion in each well correspond closely with current hot water production zones. Most feldspar separates have measured 18O values too high to have completely attained oxygen isotope exchange equilibrium with the reservoir fluid at pre-production temperatures. In general, the lower the 18O value of the feldspar, the closer the feldspar approaches exchange equilibrium with the geothermal fluid. This correlation suggests that fracture-induced increases in permeability increase both fluid

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ACCEPTED MANUSCRIPT infiltration and the surface area of the host rock exposed to geothermal fluid, promoting fluidrock interaction and oxygen isotope exchange. The two most 18O/16O-depleted feldspar samples

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have 18O values too low to be in exchange equilibrium with the pre-production reservoir fluid at

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pre-production temperatures. These discrepancies suggest that the reservoir fluid in the West

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Flank of the Coso geothermal system was hotter and/or had a lower 18O value (due to fluid-rock

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interaction at higher permeability) in the past.

1. INTRODUCTION

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The Coso Hot Springs Geothermal Field is located in China Lake, California, 10 km east of the Sierra Nevada Mountains at the western boundary of the Basin and Range Province (Fig.

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1). Coso Hot Springs is a world-class geothermal field; it is the third largest producing geothermal field in the United States, among the hottest, with a current installed power capacity

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of 273 MWe and temperatures as high as 350oC at depths as shallow as ~1 km. It is also one of the most extensively drilled geothermal systems, with many of the more than 150 wells reaching depths of 2-3 km.

The level of power production at Coso Hot Springs implies significant permeability in the reservoir rocks, which consist dominantly of metamorphosed Mesozoic diorite and younger granodiorite and granite. Although epidote, chlorite, illite, quartz, and calcite, which are typical of high-temperature geothermal systems worldwide (Henley and Ellis, 1985), are widespread at Coso Hot Springs, these minerals show no clear correlation with measured temperature in boreholes (Bishop and Bird, 1987; Lutz et al., 1996; Lutz and Moore, 1997; Kovac et al., 2004; Kovac et al., 2005). In contrast, minerals typical of moderate temperature environments, such as wairakite and prehnite, are rarely observed in the reservoir rocks.

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Figure 1: A simplified geologic map of the Coso Geothermal system (modified from Hulen, 1978) located on the Naval Air Weapons Station, China Lake, California. The West Flank of the system is defined by a North-South trending thermal high; it is the focus of this study. Line A-A’ along this thermal high is the transect for a number of cross sections used in this study. Line BB’ is a cross section line through wells 68-6 and 33A-7. Locations of wells in the West Flank used in this study are marked by red and green circles; the green circles indicate wells examined in detail for this study while the red circles indicate wells examined in previous studies. Only the wells in the West Flank with oxygen isotope data for host rocks are shown; none are shown for other areas.

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ACCEPTED MANUSCRIPT Oxygen isotope measurements provide an alternative means of mapping changes in the extent of fluid-rock interactions, independent of bulk mineralogy, and have been applied

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extensively in both active geothermal systems (e.g. Lambert and Epstein, 1980; Sturchio et al.,

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1990; Smith and Suemnicht, 1991; White and Peterson, 1991; Moore and Gunderson, 1995; McConnell et al., 1997) and fossilized hydrothermal systems (e.g. Taylor, 1971; Taylor, 1973;

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Taylor, 1974; Forester and Taylor, 1980; Criss and Taylor, 1983; Larson and Taylor, 1986a;

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Larson and Taylor, 1986b; Faure et al., 2002; Mauk and Simpson, 2007). A number of these studies of fossil meteoric hydrothermal systems have demonstrated at least a qualitative

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correlation between increased 18O/16O depletion and the abundance of hydrothermal minerals, demonstrating that the extent of 18O/16O depletion can be a guide to the extent of fluid-rock

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interaction. In their studies of the active Long Valley hydrothermal system, both Smith and Suemnicht (1991) and White and Peterson (1991) noted that well cuttings from reservoir rocks

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(typically rhyolitic tuff with initial 18Owr values of ~7.0 ‰) typically experience the greatest O/16O depletions (with 18Owr values as low as -2.6 ‰) in the deeper, higher temperature parts

of the system. Smith and Suemnicht (1991) also compared measured 18Owr values to 18Owr values calculated using then current measured temperatures, and demonstrated that many of their samples are out of oxygen isotope equilibrium with the current reservoir fluid at present temperature conditions. They concluded that this disequilibrium indicates that the Long Valley hydrothermal system has experienced protracted hydrothermal circulation at higher temperatures in the past. McConnell et al. (1997) utilized spatial patterns of 18O/16O depletion in well cuttings and core from reservoir rocks to define regional pathways of fluid circulation in the Long Valley Caldera hydrothermal system. Sturchio et al. (1990) noted the possible role of decompression boiling in producing oxygen isotope disequilibrium between current thermal waters and silica

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ACCEPTED MANUSCRIPT minerals deposited in vugs and fractures in drill core samples from the Yellowstone geothermal system. Recently, Pope et al. (2016) measured δD and δ18O values of hydrothermal epidote from

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the Krafla, Iceland geothermal system. They interpreted variations in the δD and δ18O values of

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both the epidote and reservoir fluids to result from variable mixing of shallow, sub-boiling groundwaters with vapor condensates derived from an underlying two-phase reservoir, also

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derived from local meteoric waters.

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In a previous study of the Coso geothermal system, Person et al. (2006) used δ18O values of reservoir fluids and heat flow data to constrain a regional-scale numerical model of the flow

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system supplying the Coso geothermal system. In this study we report on oxygen isotope analyses of whole-rock samples from the Coso geothermal system, and detailed whole-rock and

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mineral analyses from three producing wells. Our whole-rock results identify discrete intervals

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of maximum 18O/16O depletion in the three wells that correspond closely with zones of current

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hot water production and high permeability. These discrete, higher permeability zones have experienced significant fluid infiltration, water-rock interaction and oxygen isotopic exchange with the geothermal fluids. Our results corroborate the results of a number of previous oxygen isotope studies of (mostly) fossil hydrothermal systems, and provide a particularly clear example of the connection between the extent of 18O/16O depletion, the extent of fluid-rock interaction and permeability in crystalline rock hosts to hydrothermal systems. Therefore mapping patterns of 18O/16O depletion can be another exploration tool to define permeability structure, identify higher permeability zones, and constrain fluid flow patterns—past and present—in active hydrothermal systems, and therefore help facilitate development of enhanced geothermal systems. 1.2. Geologic Setting

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ACCEPTED MANUSCRIPT The Coso Hot Springs geothermal system is developed in Mesozoic plutonic rocks of the Sierra Nevada Batholith (Duffield et al., 1980). Diorite that was regionally metamorphosed at

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greenschist to lower amphibolite facies conditions to orthogneiss in the late Cretaceous (Bartley

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et al., 2007) (Fig. 1) is the most common rock type. Younger intrusions consist of granodiorite and granite (Duffield et al., 1980). Subsequent episodes of Cenozoic volcanism produced basalt

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and rhyolite at 4.0-2.5 Ma and 1.1-0.044 Ma (Kurliovitch et al., 2003; Duffield et al., 1980). The

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youngest episode is related to a silicic magma chamber emplaced between 5-20 km depth (Duffield et al., 1980; Reasenberg et al., 1980; Manley and Bacon, 2000). This magma body

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produces the heat driving the current geothermal activity (Duffield et al., 1980). Surface expressions of geothermal activity include opaline (SiO2) sinter and travertine

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(CaCO3) spring deposits. 14C dating of pollen in hot spring deposits has yielded ages of 11,550 to 8,550 years (Moore, unpublished data). The geothermal fluids (Table 1) have low salinities

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(5,000-10,000 ppm TDS), and are NaCl-dominated. The pre-production geothermal waters are dominated by a single-phase system with minimal boiling (Fournier et al., 1980), but there is some mineral alteration evidence for a natural clay-rich steam cap in places (Williams and McKibben, 1990). Geochemical and hydrogen isotope data indicate that the system is predominately recharged by meteoric fluids (Adams et al., 2000; Moore et al., 1990; Williams and McKibbin, 1990). The modeling results of Person et al. (2006) suggest that the geothermal fluids are recharged by lateral flow of groundwaters (meteoric waters) from fault systems and possibly the Sierra Nevada mountains to the west of the Coso field. Their model has these groundwaters infiltrating to the top of the brittle-ductile transition estimated at 4-5 km depth (Wilson et al., 2003), becoming enriched in 18O/16O to the maximum measured δ18O values of the geothermal reservoir fluids (-6 ‰) through fluid-rock isotopic exchange at elevated

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ACCEPTED MANUSCRIPT temperature, and then ascending to form the thermal plume of the Coso reservoir. Northerly trending, Basin and Range faulting and local dextral strike-slip faulting associated with the

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younger Walker Lane/Eastern California Shear Zone, guide fluid flow and influence current

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geothermal activity (Monastero et al., 2005; Davatzes and Hickman, 2010; Kaven et al., 2012). Further, Davatzes and Hickman (2010) propose that stress-induced fracture slip leads to fluid

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flow and associated mineral alteration, which suggests that permeability will dynamically evolve

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as this system evolves over time. Additional details about the geology, including hydrothermal alteration, are provided by Hulen (1978), Duffield et al. (1980), Lutz et al. (1996), Adams et al.

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(2000), Manley and Bacon (2000), Kovac et al. (2005) and Etzel (2016).

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Table 1: The composition of pre-production geothermal fluids from wells 68-6, 33A-7 and 7319. Chemistry in ppm by weight; 18O and D in per mil (‰) relative to VSMOW. 73-19 2044 579 50 27 N/A 89 602 44 16 3783 2 -6.0 ‰ -96 ‰

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33A-7 1273 187 7 11 N/A 58 514 124 40 2163 3 -7.5 ‰ N/A

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Na K Ca Li Mg B SiO2 HCO3 SO4 Cl F 18  OW D Date Measured

68-6 748 136 7 7 <1 39 596 102 30 1260 3 -8.5 ‰ -104 ‰

1.3. Thermal Structure Borehole measurements have defined a north-south trending zone of high temperatures on the West Flank of the field (Fig. 1). Geochemical and thermal data suggest that the system is compartmentalized with upflow zones in the northeast and southwest (Williams and McKibben,

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ACCEPTED MANUSCRIPT 1990). Figure 2 shows the thermal structure of the West Flank of the system based on preproduction temperature data. Maximum measured temperatures at depth reach 350oC. Reservoir

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temperatures in the south are higher than in the north at equivalent elevations. The thermal

northward.

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2. METHODS

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pattern defines a plume of hot water that ascends in the south and then migrates laterally

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Rock cuttings were sampled at regular intervals (nominally ~30 - 80 m apart) in wells 686, 33A-7 and 73-19 for oxygen isotope, X-ray diffraction (XRD) and petrographic analyses. Thin

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sections were examined using a petrographic microscope under transmitted light to identify primary and secondary (alteration) minerals, vein mineral assemblages and their paragenesis,

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structural features and the host rock lithology. Etzel (2016) provides a detailed presentation of these results.

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Clay mineralogy and clay mineral abundances have been determined by x-ray diffraction analyses of whole-rock and clay-size fractions (<5 m) at the Energy & Geosciences Institute at the University of Utah, using a Bruker D8 Advance x-ray diffractometer. The clay mineralogy was determined using methods described by Moore and Reynolds (1997) on air-dried and glycolated scans of the clay-sized fraction. Phase quantification using the Rietveld method was performed using TOPAS software, developed by Bruker AXS, as described in Lutz and Moore (1997). Mineral abundances are in weight percent with values rounded to the nearest whole number. Values of less than one weight percent obtained from Rietveld analysis are reported as trace amounts. Additional details of these techniques, clay mineralogy and of host rock lithology are reported in Etzel (2016).

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Figure 2: General north-south thermal structure of the West Flank along the transects A-A’ and B-B’. Temperature values come from down-hole measurements made prior to production. Only wells 68-6, 33A-7 and 73-19 are shown for reference; in total, thermal data from 25 wells in the West Flank were used to construct these cross-sections. Measurements are available to elevations approaching -2km in almost all of these wells; therefore, these temperature contours are well constrained at all depths displayed.

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ACCEPTED MANUSCRIPT Oxygen isotopic measurements were conducted at the University of Wisconsin-Madison on 91 whole rock samples and 30 feldspar separates from wells 68-6, 33A-7 and 73-19. Mineral

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separates were made using a Frantz magnetic separator followed by hand-picking under a

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stereoscopic microscope. The oxygen isotope analyses were done by laser-aided fluorination using a lasing sample chamber outfitted with an airlock sample chamber to prevent pre-

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fluorination of reactive rock powders (Spicuzza et al., 1998). Oxygen isotope values are reported

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in the standard 18O notation, relative to VSMW, and are standardized using UWG-2 garnet (18O = 5.8 ‰; Valley et al., 1995). Multiple analyses of UWG-2 (n>4) within individual

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analytical sessions yield a reproducibility of ± 0.07 ‰ (2The average measured δ18O values

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for all analyses of UWG-2 done in this study (n = 34) is 5.71 ± 0.20 ‰ (2, within error of the

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published value for this standard. The 513 whole rock oxygen isotope analyses made available by Terra-Gen were analyzed at Southern Methodist University. These oxygen isotope analyses

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were determined using the bromine pentafluoride extraction procedure described by Clayton and Mayeda (1963). The δ18O analyses are reported relative to VSMOW, and long-term analyses of in-house standards indicate reproducibility of ± 0.2‰. Isotopic measurements of the geothermal fluid (Table 1) were also made at Southern Methodist University and are reported relative to VSMOW. For oxygen, isotopic compositions were determined following a method similar to that of Epstein and Mayeda (1953) and hydrogen oxygen isotope compositions were determined following a method similar to what was described by Bigeleisen et al. (1952). In-house standards indicate a long term reproducibility of ± 0.1 ‰ for 18O and ± 1 ‰ for D. 3. RESULTS 3.1. Host Rock and Alteration Mineralogy

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ACCEPTED MANUSCRIPT The reservoir rocks observed in wells 68-6, 33A-7 and 73-19, the three wells studied in detail (well locations shown in Fig. 1) are principally diorite orthogneiss with subordinate diorite

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and granodiorite.

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Two episodes of overprinting on the primary (igneous) mineralogy of the reservoir rocks are recognizable (Bishop and Bird, 1987; Lutz et al., 1996; Lutz and Moore, 1997; Kovac et al.,

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2005; Etzel, 2016). Regional lower amphibolite to greenschist facies metamorphism (late

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Cretaceous) has produced plagioclase, hornblende, biotite, epidote, titanite, chlorite and illite (muscovite) throughout all three wells (figs. 3, 4a, b). During subsequent geothermal activity,

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chlorite and rare interlayered smectite/chlorite replaced hornblende and biotite, and epidote, illite (muscovite) and other clay minerals replaced plagioclase (Fig. 4c-f). This geothermal activity has

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resulted in progressive changes in temperature sensitive clay minerals with depth. Throughout the field, smectite (S) is present at the shallowest depths, followed by interlayered illite/smectite

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(I/S) and rare chlorite/smectite, and finally illite and chlorite at deeper levels. The deepest occurrence of smectite above trace levels (> 1 wt %) defines the base of the smectite zone at an elevation of about 1000 meters in wells 68-6 and 33A-7 (Fig. 3). In well 7319, smectite occurs only at trace levels (< 1 wt %) and occurs persistently only at elevations above ~300 m (Fig. 3). The deepest occurrence of interlayered illite-smectite above trace levels defines the base of the I/S zone in wells 33A-7 and 73-19 at elevations of -41 meters and -143 meters, respectively. Interlayered illite-smectite is completely absent in well 68-6. Trace amounts of the lower temperature smectite and interlayered I/S clays exist within the illite zone in 33A-7 (Fig. 3).

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Figure 3: Clay mineral content and measured pre-production temperature profiles for wells 68-6, 33A-7 and 73-19. I/S = interlayered illite-smectite; C/S = interlayered chlorite-smectite. Dashed lines denote locations at which clay transitions occur; measured temperatures at these locations are in parentheses. Boiling point curves calculated following the approach presented by Haas (1971).

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Figure 4: Photomicrographs of the lithologies and alterations observed in the Coso geothermal system. Panels a) and b) show a typical unaltered meta-diorite sample with hornblende (Hbl), untwined feldspar (Fsp), and rare anhedral titantite (Ttn). This sample is from 33A-7_1100. Panels c) and d) show extensive geothermal-related alteration in feldspar, hornblende and biotite found in 68-6_2792. Feldspar is altered to illite (Ill); hornblende and biotite are completely altered to chlorite (Chl). Panels e) and f) show a feldspar partially altered to illite found in 686_2941. The field of view in panels a) – d) is 1.6 mm and in panels e) and f) is 0.8 mm. The transitions from smectite to interlayered illite/smectite to illite typically occur at temperatures of ~180oC and 225oC, respectively (Steiner, 1968; Browne, 1978; Henley and Ellis, 1983; Browne, 1984; Reyes, 1990). However, in the three Coso wells, these transitions do not always occur at or near these temperatures, based on measured pre-production temperatures in

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ACCEPTED MANUSCRIPT wells (Fig. 3). In 33A-7, the smectite to interlayered illite/smectite and illite/smectite to illite transitions occur well below expected temperatures at ~123oC and ~210oC, respectively. In well

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73-19, the smectite to interlayered illite/smectite transition occurs near the expected T of 180oC, but interlayered illite/smectite persists to depths where temperatures exceed 280oC. In 68-6

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smectite disappears near a pre-production temperature of 180oC (at an elevation of ~1000

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meters), its expected limit of thermal stability, but no interlayered illite/smectite is present below

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the smectite. These discrepancies suggest that temperature variations alone cannot explain the detailed distribution of clay minerals observed in these wells.

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3.2. Whole Rock Oxygen Isotope Characteristics of the Coso Hot Springs Geothermal System General patterns of 18O variation as a function of elevation in the West Flank of the

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Coso geothermal field are illustrated in figure 5. These plan contour maps are based on 513

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whole rock 18O analyses of well-cuttings supplied by Terra-Gen Operating Company

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(Supplementary Data Table A.1), plus 91 additional whole rock 18O values measured in this study (Table 2). The plan maps of 18O variations shown in figure 5 were constructed by averaging the 18O values within 480 m intervals in each well; sufficient analyses were not available deeper than 1125 m below sea-level to construct a useful plan map. Contouring was done using the MATLAB contour function. This function requires position (in this case we used northing and easting values recorded during drilling) and a calculated average 18O value within each interval for each well used. Average 18O values used for contouring are presented in Supplementary Data Table A.3. Near surface 18O values (surface to 800 m elevation) are similar with, or slightly lower than, measured 18O values for primary igneous rock in the Central Sierra Nevada Batholith (+7.5 ‰ to +9.0 ‰; Masi et al., 1981; Lackey et al., 2008) except for depletion to 18O values between +5 ‰ and +6 ‰ in the northwest quadrant of the system.

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Figure 5: Contoured plan maps of the whole rock oxygen isotope data at a series of elevation intervals in the Coso system. The area of each slice is equal to the area of the West Flank outlined in Figure 1 but the horizontal dimension has been exaggerated for visual clarity. On each map, the 18O values for each well location are the average of all 18O data from that well within the 480 m elevation interval represented by that map. The contour values are in ‰. Dashed lines are used where contouring has been inferred. All elevations are relative to sea level. The green and black dots are well locations; green dots are wells 68-6, 33A-7 and 73-19. The location of transect A-A’ in the West Flank is shown in the upper left panel. Contours were generated using the MATLAB contour function.

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ACCEPTED MANUSCRIPT With increasing depth, significant and widespread decreases in 18O occur only in the northwest quadrant, where 18O values can decrease down to values between 0 and +1 ‰ (-645 to -1125 m

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elevation interval). Regardless of depth, the host rock in most of the southern half of the

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reservoir does not experience significant 18O/16O depletion below +5 ‰. At the sampling scale of figure 5, the thermal upflow zones inferred in the SW and NE sectors of the field by Williams

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and McKibben (1990) are not reflected in significant lowering of 18O values in these sectors.

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Figure 6 illustrates variations in whole rock 18O values within the vertical section A-A’ shown in figure 1 (along the West Flank). Again, contouring with a MATLAB contour function,

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but using all whole rock 18O values (no averages). In general, depletions in 18O/16O are more

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extensive and extend to significantly greater depth in the northern portion of the cross-section. In

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detail, localized intervals of significant 18O/16O depletion, with less depleted rocks above and below, occur in the northern wells 68-6 and 33A-7 in intervals centered on elevations of 200, -

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200 and -1200 m. Maximum depletions to 18O values below 0 ‰ occur in the lowest of these intervals. In contrast, host rocks in the southern part of the traverse are not significantly depleted except for two or three discrete intervals in 73-19 and 58A-18. These spatial patterns of localized and significant 18O depletions do not reflect the spatial 18O patterns expected from the low and smooth temperature gradients (dT/dx) characteristic of regional metamorphism. The observed spatial patterns of 18O variation are instead more consistent with localized 18O depletions expected from fluid-rock interactions related to discrete zones of high permeability associated with geothermal activity. The thermal structure along the cross-section A-A’ (Fig. 2) is also superimposed on this 18O cross-section in figure 6 which shows that variations in pre-production temperatures are not, in general, well correlated with variations in whole-rock 18O values. With the few localized

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Figure 6: Vertical cross section of the whole rock oxygen isotope data along the north-south transect, A-A’. The locations of the nine wells on or close to the A-A’ traverse (Fig. 1), and data points from each well used for contouring δ18O values are shown; contours are in ‰. Red circles indicate the locations of samples analyzed in this study. Isotherms based on preproduction downhole measurements from figure 2 are superimposed to define areas along the cross section above 300 oC (dark gray) and between 200 oC and 300 oC in varying shades of gray; see text for discussion. The major production zone in each well is noted.

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ACCEPTED MANUSCRIPT exceptions noted above, the host rocks are not depleted in 18O/16O to any significant extent in the southern part of the traverse, even within the deeper, higher temperature levels. In the northern

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part of the section, hotter host rocks below -1400 m are less depleted in 18O/16O than cooler rocks immediately above between -1000 and -1400 m. However, depth intervals of current production

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in all three wells do correlate with zones of maximum 18O/16O depletion (Fig. 6). This correlation

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suggests that permeability plays an important role in controlling the extent of fluid-rock

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interaction—and hence the extent of 18O/16O depletion—experienced by the host rocks in the Coso geothermal system.

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3.3. Oxygen Isotope Characteristics of the West Flank 3.3.1. Whole Rock 18O Analyses. Wells 68-6, 33A-7 and 73-19 were selected for detailed

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oxygen isotope analyses of whole-rock and mineral separates (table 2; Figure 7). In well 68-6

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(Fig. 7a) whole-rock δ18O values are somewhat to moderately depleted in 18O/16O relative to

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primary igneous values (+7.5 to 9.0‰), and fluctuate between 1.8 ‰ and +6.0 ‰ down to ~-700 meters elevation, with a tendency to decrease with depth. Below -700 m, 18O/16O depletions are even greater, and the maximum δ18O value is +3.0 ‰, with most δ18O values below +1 ‰. Whole rock δ18O values reach a minimum of -4.6 ‰ at -1552 meters above sea level (msl) (Fig. 7a). In well 33A-7 δ18O values generally decrease with elevation from a maximum value of +7.6 ‰ at an elevation of 1251 meters to a minimum of -3.1 ‰ at -1195 meters, just above the production zone (interval of lost circulation) in the well (Fig. 7b). Immediately below, and to the bottom of the well, δ18O values are significantly higher, ranging between +3 to +4.9 ‰, but are still significantly lower than primary igneous values. The host rocks in well 73-19 are less depleted in 18O/16O compared to the rocks of the other two wells (Fig. 7c). The highest measured δ18O values from whole rock samples overlap primary igneous 18O values of +7.5 to 9.0 ‰

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ACCEPTED MANUSCRIPT Table 2: Measured δ18O values for whole rock and feldspar mineral samples from well 68-6, 33A-7 and 73-19, Coso geothermal system, California. Values are reported in per mil (‰) notation, relative to VSMOW. Elevation is in meters above sea level (msl).

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Well: 33A-7 Well: 73-19 WholeWholeRock Feldspar Elevation Rock Feldspar Elevation (m) (‰) (‰) (m) (‰) (‰) 1251 7.62 7.50 996 6.98 7.79 1175 3.70 886 6.61 1023 5.39 877 6.37 870 7.45 822 7.48 8.12 718 4.84 722 5.52 639 5.34 490 7.57 413 0.20 469 6.36 261 3.59 207 7.16 185 3.00 5.49 195 7.41 -41 4.18 42 6.07 -283 3.89 -49 6.63 -360 2.71 -79 5.91 -585 1.03 -119 4.61 6.14 -689 3.93 -143 2.38 -833 1.60 -213 6.30 7.26 -904 -0.44 -313 5.94 -1005 -0.02 -395 6.85 -1005 -1.02 -441 5.13 -1031 -0.98 -544 5.14 6.20 -1059 0.14 -565 3.30 6.34 -1114 0.51 -2.37 -586 3.30 -1170 -2.35 -1195 -3.08 -0.03 +0.63* -1230 0.94 4.16 -1284 2.15 4.20 -1314 1.55 -1343 2.92 -1230 2.10 4.60 -1402 3.91 -1487 4.48 5.49 -1603 4.60 -1632 4.25 -1661 4.05 -1690 1.78 -1721 4.21 -1751 4.03 -1782 4.61 -1812 4.10 -1873 4.49 -1934 3.88 5.90 # Hand picked cloudy (more altered) feldspar * Feldspar separated from finer-grained (150-200 mesh) whole rock aliquot

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Well: 68-6 WholeElevation Rock Feldspar (m) (‰) (‰) 996 3.88 937 4.47 6.04 855 2.55 789 3.49 6.59 710 5.08 587 5.78 556 6.02 6.29 394 3.57 293 4.07 208 4.30 31 4.43 -60 2.44 5.24 -129 2.87 -165 1.50 4.96 -230 2.60 -365 3.78 5.64 -455 4.67 6.76 -581 4.46 -707 2.94 -911 2.70 -1020 0.75 -1091 0.75 3.09 -1188 0.61 -1327 0.75 3.08 -1407 2.83 3.59 -1463 0.54 -1511 0.26 -1552 -4.60 -3.88 -5.06# -4.12* -1594 -1.05 1.76 -1706 0.64

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Figure 7: Measured 18O values of whole rock and feldspar samples as a function of elevation for well 68-6 (panel a), 33A-7 (panel b) and 73-19 (panel c). All values are reported in per mil (VSMOW), and plotted at elevation (meters above sea level). The shaded light blue regions represent the range of primary 18O values of diorite to granodiorite rocks from the region that are equivalent to the reservoir host rocks in the Coso system (Lackey et al., 2008). The dotted line marks the current production zone (interval of lost circulation) in each well. (Masi et al., 1981; Lackey et al., 2008). The only δ18O values below +5 ‰ occur at elevations of -119 m and below. The minimum δ18O value in well 73-19 is still positive, +2.4 ‰ (-143 meters elevation, Fig. 7). In all three wells, samples with the lowest measured δ18O values (or with

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ACCEPTED MANUSCRIPT 18Owr ≤ 2.0 ‰, well 33A-7) correspond with major intervals of lost circulation, the primary production intervals for each well (Fig. 7).

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3.3.2. Feldspar vs. Whole Rock Oxygen Isotope Compositions. Measured 18O in bulk feldspar

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separates range from 6.8‰ to -3.9 ‰ in well 68-6 and from 7.5 ‰ to -2.4 ‰ in 33A-7; feldspars

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from well 73-19 have higher 18O values ranging from 8.1 ‰ to 6.1 ‰. In all three wells, the 18O values of the feldspar correlate positively with the whole-rock 18O values, and are

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systematically 1 to 3 ‰ higher than the whole rock values, even in the more 18O/16O -depleted

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samples (Fig. 7a). The only exception is in 33A-7 at -1139 msl. The bulk feldspar separates are mixtures of primary (igneous or metamorphic) feldspar and feldspar that has experienced oxygen isotope exchange with geothermal fluid (hydrothermal feldspar). The hydrothermal feldspar

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should have a considerably lower 18O value than primary feldspar, owing to the low 18O value of the geothermal fluid and intermediate temperature of exchange (<350oC). Further, feldspar is

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normally more susceptible to oxygen isotope exchange compared to hornblende and less abundant quartz in the host diorite and quartz diorite (Taylor and Forester, 1979; Criss and Taylor, 1983; Cole et al., 1992). As a result, hydrothermal feldspar might be expected to have a 18O value less than that of the whole-rock for rocks that are only moderately depleted in 18

O/16O. In thin section, primary igneous and regional metamorphic feldspar are clear and

twinned (igneous feldspar); hydrothermal feldspar is cloudy and sometimes turbid from the presence of pits/cavities and very small grains of clay minerals. The intergrown nature of primary and hydrothermal feldspar at the grain scale makes their effective separation by standard density means impractical. However, a concentrate of cloudy-looking feldspar was made by hand picking the bulk feldspar separate from sample 68-6_-1552. Analysis of this concentrate yielded a 18O value of -5.1 ‰, significantly lower than the bulk feldspar (-3.9 ‰) and whole rock (-4.6

21

ACCEPTED MANUSCRIPT ‰) 18O values of this sample (Table 2). In-situ (SIMS) analysis will be necessary to define clearly the 18O value of hydrothermal feldspar in the host rock (e.g., Valley and Graham, 1996;

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Valley and Kita, 2009). Regardless, this systematic difference between measured 18O values of

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bulk feldspar and whole rock suggests that the 18O values of both feldspar and whole rock

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reflect the extent of interaction (and oxygen isotopic exchange) between host rock and geothermal fluid.

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Another possibility for the systematic 18O/16O enrichment in bulk feldspar relative to the

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whole rock is that the lower 18O, hydrothermally altered feldspar was preferentially removed by crushing during preparation of the sample aliquots (>80 or >100 mesh size) from the whole rock

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used to make plagioclase separates. To test this possibility, additional analyses were made of

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feldspar separated out of the finer-grained (150 to 200 mesh) residual fraction of the whole rock aliquots for two samples. The analyzed 18O values for these finer-grained feldspar separates are

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-4.1 ‰ (compared to -3.9) for sample 68-6_-1551, and +0.6 ‰ (compared to -0.0 ‰) for sample 33A-7_-1195 (Table 2). Additional details of these results are in Etzel (2016). Feldspar separates from these finer-grained fractions are not significantly lower in 18O compared to feldspar separates from the coarser-grained fractions. The systematically higher measured 18O values of feldspar relative to the whole rock indicate, from mass balance considerations, that there is at least one other major phase (likely chlorite and/or other clay mineral alteration products) in the rock that is significantly depleted in 18

O/16O relative to the whole rock value. Chlorite can replace hornblende but would be stable

with biotite during the greenschist facies conditions of the earlier regional metamorphism; this generation of chlorite will not be related to geothermal activity. Replacement of biotite by chlorite would more likely occur at the sub-greenschist facies conditions of the geothermal

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ACCEPTED MANUSCRIPT environment. Hence, bulk separates of chlorite will potentially contain two generations of chlorite of very different 18O value. Good separation of bulk samples of chlorite by magnetic

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and density methods proved impractical owing to the intergrowth of chlorite with biotite and

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hornblende at the grain size scale. Concentrates of chloritized biotite from four samples were

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handpicked. Their 18O values, compared to whole rock values (in parentheses) are: 1) 0.2 ‰ (3.8), sample 68-6_-382; 2) -3.2 ‰ (0.7), sample 68-6_-911; 3) -3.8 ‰ (-3.1), sample 33A-7_-

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1195; and 4) -0.45 ‰ (4.6), sample 73-19_-119. These chloritized biotites are from almost 1 to as much as 5 ‰ depleted in 18O/16O relative to the whole rock values, satisfying at least

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qualitatively this mass balance requirement. In-situ analyzes will be necessary to define the 18O

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values of the hydrothermal chlorite more precisely.

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4. DISCUSSION

4.1. Clay Mineralogy as a Function of Fluid-Rock Interaction

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The observed distributions of clay minerals in wells 68-6, 33A-7 and 73-19 are inconsistent with the pre-production temperatures in these wells (Fig. 3). These discrepancies may indicate that the pre-production thermal regime was different than the regime responsible for clay mineral zoning, and that the clay minerals have not yet re-adjusted to the newer (preproduction) thermal conditions. The persistence of clay minerals outside their normal stability fields (e.g., smectites at T> 180 oC in in 73-19, Fig. 3) is common but the reasons are poorly understood. Nucleation and reaction kinetics may play roles. Permeability, which can either promote (high permeability) or inhibit (low permeability) clay-fluid interactions, may also be influencing the spatial distribution of the clay minerals. Oxygen isotope evidence suggests that the extent of fluid-rock interaction experienced by the host rock is influencing the distribution of clay minerals in the Coso system. The amount of

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O/16O depletion is a monitor of the extent of fluid-rock interaction by the host rock—and hence

it’s permeability. Figure 8 superimposes the distribution of clay minerals in wells 68-6, 33A-7

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and 73-19 on variations in whole rock 18O values along the cross section A-A’. The host rocks

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are systematically more depleted in 18O/16O at depth at the north end compared to the south end of the traverse. In addition, significant 18O/16O depletions extend to markedly shallower depth at

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the north end of A-A’. The overall extent of 18O/16O depletion in the south third of the traverse is

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considerably less, and significant 18O/16O depletions in well 73-19 are confined to several discrete/narrow/short intervals. Hence in well 73-19 the transition from smectite to interlayered

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illite/smectite clay is better correlated with temperature, as this transition occurs near the expected temperature of 180oC (Fig. 3). However, interlayered illite/smectite clay occurs as

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deeply as -143 m elevation (T = 282oC, Fig. 3) in well 73-19, an elevation corresponding to a

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local interval of significant 18O/16O depletion (δ18O < 4 ‰). Figure 6 shows that there is a

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general correlation between greater and more spatially widespread 18O/16O depletion in the north end of A-A’ with a thinner smectite zone in 68-6 and 33A-7 and the lack of a mixed illite/smectite transition zone in 68-6. Additionally, the transition from interlayered illite/smectite to illite clay in 73-19 is correlated with local 18O/16O depletion. These highly 18O/16O depleted zones within traverse A-A’ record greater fluid-rock interaction, and are characterized by higher permeability based upon the greater extent of 18O/16O depletion relative to other portions of the reservoir. Hence the extent of fluid-rock interaction (host rock permeability) appears to be controlling, at least in part, the spatial distribution of clay minerals and clay zoning patterns in wells 68-6, 33A-7 and 73-19 along the traverse A-A’.

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Figure 8: Clay zones within wells 68-6, 33A-7 and 73-19 superimposed onto the contoured cross-section of 18O values along the traverse A-A’ in the West Flank. Black circles represent intervals with preexisting 18O values, red circles represent intervals where 18O values were determined for this study. Increasing extent of 18O/16O depletion northward correlates with changes in clay mineral zoning with depth, suggesting a correlation between clay mineral stability, extent of fluid-rock interaction and permeability (see text for discussion) 25

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4.2. Oxygen Isotope Exchange and Fluid-Rock Interaction

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The δ18O values of the whole rock samples from the West Flank range widely from -4.6 ‰ to 7.6 ‰ (Table 2). Reservoir rocks in 33A-7 and 68-6, except for the shallowest levels, have

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experienced at least moderate 18O/16O depletions (δ18O <5.0‰) from primary igneous δ18O

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values (+7.5 to 9.0‰; Lackey et al., 2008). Hence most of the reservoir rock sampled in these

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two wells has experienced at least some interaction and oxygen isotope exchange with the geothermal reservoir fluid derived from low δ18O local meteoric water. Further, a few discrete

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intervals in 33A-7 and 68-6 have experienced significantly greater 18O/16O depletion, and have negative whole-rock δ18O values ranging as low as -4.6 ‰.

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Significant variations in the extent of 18O/16O depletion have been documented in other active hydrothermal systems, such as the system at Long Valley, CA, where in general, rock

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samples from deeper and higher T sections of wells tend to have lower δ18O values (Smith and Suemnicht, 1991; White and Peterson, 1991). Both these studies and McConnell et al. (1997) attribute these large-scale variations in δ18O values primarily to variations in temperature—both spatially and over time. In the West Flank at Coso, the δ18O values generally decrease in well 33A-7 with increasing depth (and increasing T) down to the production zone (Fig. 7), so temperature is likely playing a role in determining the extent of 18O/16O depletion at large spatial scales in the Coso system as well. These previous studies at Long Valley also document smallerscale variations in the extent of 18O/16O depletion that they attribute to incomplete exchange (either from kinetic barriers or limited permeability), or to varying abundances of less easily and more easily altered minerals/rock components in the samples analyzed (McConnell et al., 1997). In the three wells analyzed at Coso in this study, localized intervals of significantly greater

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O/16O depletion than in rocks immediately above or below (Fig. 7) indicate that these reservoir

host rocks have experienced significant variations in the extent of fluid-rock interaction.

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The extent of 18O/16O depletion is a function of the amount of reservoir fluid with which

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the rock has interacted, the temperature and the degree to which isotope exchange equilibrium is attained. Box model water (W) to rock (R) ratios (Taylor, 1971) have been used routinely to

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estimate the amounts of water involved in fluid-rock interaction in both fossil and active

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hydrothermal systems. Studies that have developed reactive transport models applied to oxygen isotopes (Baumgartner and Rumble, 1988; Bowman and Willett, 1991; Bowman et al., 1994;

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Cook et al., 1997; Baumgartner and Valley, 2001) point out that there are several assumptions involved and resulting limitations in applying W/R ratios to isotopically exchanged rocks within

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hydrothermal systems. In particular, the impact of the prior exchange history of the fluid along flow path(s) to local sites of hydrothermal alteration and isotopic exchange will mean that box

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model W/R ratios will underestimate actual W/R ratios by an amount that increases with increasing distance along the reactive flow path. The actual amount of fluid involved might have been orders of magnitude higher. As a result, calculated box model W/R ratios should be regarded as minimum estimates of the amounts of fluid involved in hydrothermal alteration and isotopic exchange at specific sites within hydrothermal flow systems. Meteoric waters recharging the Coso geothermal reservoir have experienced considerable 18O/16O enrichment, likely resulting from fluid-rock exchange within the deeper, hotter portions of the Coso hydrologic system (Person et al., 2006). However, the geothermal fluids are still well out of oxygen isotope exchange equilibrium with the reservoir host rocks at temperatures from 200 to 350oC. As a result, these fluids are still chemically (isotopically) reactive, and are capable of producing significant 18O/16O depletions in the infiltrated rocks. Consequently, variations in the extent of

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O/16O depletion in rocks from local segments of the Coso hydrothermal system—and variations

in calculated W/R ratios—will qualitatively reflect variations in fluid fluxes. Hence, the W/R

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ratios calculated from oxygen isotope ratios are expected to reflect the relative differences in

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permeability and time-integrated fluid flux for these local segments.

Calculated W/R ratios (atomic oxygen) using both closed and open system box models

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(Taylor, 1971; Criss and Taylor, 1986) are presented in Tables 3-5. Procedures and data used for

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these calculations are presented in Appendix B. The δ18O measurements identify a limited number of localized intervals of much more extensive 18O/16O depletion within the Coso

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reservoir host rocks that have interacted and exchanged isotopically with significantly larger quantities of geothermal fluid (higher calculated W/R ratios). Well 33A-7 has three such

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intervals of locally more extensive 18O/16O depletion and higher calculated W/R ratios at -1170 -1195 msl [(W/R)c = 1.94]; at 413 msl [(W/R)c = 1.35]; and at 1175 msl [(W/R)c = 0.71]. Well

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68-6 has one interval at -1552 msl that is characterized by much greater 18O/16O depletion and much higher W/R ratio [(W/R)c = 2.45] than elsewhere in the well. In general, calculated W/R ratios for 73-19 are lower than for the other two wells, suggesting that permeability in this well is generally lower than in 68-6 and 33A-7. However, locally higher W/R ratios (closed system) are computed for elevations of -586 (0.57), -565 (0.57), -143 (0.70) and 722 (0.28) meters in well 73-19. The presence of these limited number of discrete intervals of significant 18O/16O depletion (and high calculated W/R ratios) in the three wells reflect discrete zones of much higher permeability within low permeability (much less 18O/16O depleted) rock. This spatial pattern is

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Table 3: Calculated equilibrium r-w and fsp-w, measured r-w and fsp-w, and W/R ratios for well 68-6. W.R. = whole-rock; Eq = equilibrium; FSP = feldspar. 68-6 W.R. Eq Measured FSP Eq Measured W/R W/R 18 1 2 18 1 2 ID closed Open  O r-w r-w  O fsp-w fsp-w 3.88 996 9.33 12.55 0.49 0.40 937 4.47 9.33 13.14 6.04 10.17 14.71 0.36 0.31 855 2.55 9.33 11.22 0.80 0.59 789 3.49 9.27 12.16 6.59 10.10 15.26 0.58 0.46 5.08 710 9.27 13.75 0.22 0.20 587 5.78 9.03 14.45 0.06 0.05 556 6.02 8.79 14.69 6.29 9.59 14.96 0.00 0.00 394 3.57 8.04 12.24 0.57 0.45 4.07 293 7.68 12.74 0.45 0.37 208 4.30 7.36 12.97 0.40 0.33 31 4.43 7.36 13.10 0.37 0.31 -60 2.44 6.81 11.11 5.24 7.46 13.91 0.83 0.60 2.87 -129 6.81 11.54 0.73 0.55 -165 1.50 6.81 10.17 4.96 7.46 13.63 1.04 0.71 -230 2.60 6.81 11.27 0.79 0.58 -365 3.78 6.72 12.45 5.64 7.37 14.31 0.52 0.42 -455 4.67 6.51 13.34 6.76 7.13 15.43 0.31 0.27 -581 4.46 6.38 13.13 0.36 0.31 -707 2.94 6.30 11.61 0.71 0.54 -911 2.70 6.06 11.37 0.77 0.57 -1020 0.75 5.94 9.42 1.22 0.80 -1091 0.75 5.88 9.42 3.09 6.46 11.76 1.22 0.80 -1188 0.61 5.79 9.28 1.25 0.81 -1327 0.75 5.50 9.42 3.08 6.05 11.75 1.22 0.80 2.83 -1407 5.60 11.50 3.59 6.16 12.26 0.74 0.55 -1463 0.54 5.77 9.21 1.27 0.82 -1511 0.26 5.98 8.93 1.33 0.85 -1552 -4.60 5.98 4.07 -3.88 6.57 4.79 2.45 1.24 -1.05 -1594 5.73 7.62 1.76 6.30 10.43 1.63 0.97 -1706 0.64 5.00 9.31 1.24 0.81 Equilibrium r-w and fsp-w fractionation values calculated at pre-production reservoir temperatures using the experimental fractionation factor for An50-water and An25-water (O’Neil and Taylor, 1967), respectively. 2 Measured wr-w and fsp-w are the differences between the measured 18O values of whole-rock or feldspar, and the measured 18O value of the current geothermal fluid (-8.67 ‰). 1

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ACCEPTED MANUSCRIPT Table 4: Calculated equilibrium r-w and fsp-w, measured r-w and fsp-w, and W/R ratios for well 33A-7. W.R. = whole-rock; Eq = equilibrium; FSP = feldspar. W.R. 18O

Eq 1 r-w

Measured 2 r-w

FSP 18O

1 fsp-w

Measured 2 fsp-w

W/R closed

W/R Open

1251 1175 1023 870 718 639 413 261 185 -41 -283 -360 -585 -689 -833 -904 -1005 -1005 -1031 -1059 -1138 -1170 -1195 -1230 -1276 -1284 -1314 -1343 -1402 -1487 -1603

7.62 3.7 5.39 7.45 4.84 5.34 0.2 3.59 3 4.18 3.89 2.71 1.03 3.93 1.6 -0.44 -0.02 -1.02 -0.98 0.14 0.51 -2.35 -3.08 0.94 2.15 1.55 2.92 2.1 3.91 4.48 4.6

23.44 19.54 14.03 11.00 11.00 9.78 9.30 8.70 8.56 8.20 7.94 7.76 7.39 7.20 6.81 6.70 6.49 6.49 6.44 6.38 6.18 6.06 6.02 6.02 5.96 5.85 5.79 5.73 5.60 5.48 5.25

15.11 11.19 12.88 14.94 12.33 12.83 7.69 11.08 10.49 11.67 11.38 10.20 8.52 11.42 9.09 7.05 7.47 6.47 6.51 7.63 8.00 5.14 4.41 8.43 9.64 9.04 10.41 9.59 11.40 11.97 12.09

7.50

25.34

14.99

5.49

9.34

0.00 0.71 0.40 0.03 0.50 0.41 1.35 0.73 0.84 0.62 0.68 0.89 1.20 0.67 1.09 1.46 1.39 1.57 1.56 1.36 1.29 1.81 1.94 1.21 0.99 1.10 0.85 1.00 0.67 0.57 0.55

0.00 0.54 0.34 0.03 0.41 0.35 0.85 0.55 0.61 0.49 0.52 0.64 0.79 0.51 0.74 0.90 0.87 0.94 0.94 0.86 0.83 1.03 1.08 0.79 0.69 0.74 0.62 0.69 0.51 0.45 0.44

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Eq

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33A-7 ID

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12.98

-2.37

6.78

5.12

-0.03 4.16 4.20

6.61 6.61 6.55

7.46 11.65 11.69

4.60

6.30

12.09

5.49

6.03

12.98

Equilibrium r-w and fsp-w fractionation values calculated at pre-production reservoir temperatures using the experimental fractionation factor for An50-water and An25-water (O’Neil and Taylor, 1967), respectively. 2 Measured wr-w and fsp-w are the differences between the measured 18O values of whole-rock or feldspar, and the measured 18O value of the current geothermal fluid (-7.49 ‰). 1

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ACCEPTED MANUSCRIPT Table 4 (continued): Calculated equilibrium r-w and fsp-w, measured r-w and fsp-w, and W/R ratios for well 33A-7. W.R. = whole-rock; Eq = equilibrium; FSP = feldspar.

-1632 -1661 -1690 -1721 -1751 -1782 -1812 -1873 -1934

4.25 4.05 1.78 4.21 4.03 4.61 4.1 4.49 3.88

5.20 5.16 5.13 5.13 5.00 4.95 4.84 4.76 4.68

11.74 11.54 9.27 11.70 11.52 12.10 11.59 11.98 11.37

FSP 18O

Eq 1 fsp-w

Measured 2 fsp-w

W/R closed

W/R Open

0.61 0.65 1.06 0.62 0.65 0.55 0.64 0.57 0.68

0.48 0.50 0.72 0.48 0.50 0.44 0.49 0.45 0.52

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Measured 2 r-w

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Eq 1 r-w

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W.R. 18O

5.90

5.17

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33A-7 ID

13.39

Equilibrium r-w and fsp-w fractionation values calculated at pre-production reservoir temperatures using the experimental fractionation factor for An50-water and An25-water (O’Neil and Taylor, 1967), respectively. 2 Measured wr-w and fsp-w are the differences between the measured 18O values of whole rock or feldspar, and the measured 18O value of the current geothermal fluid (-7.49 ‰).

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ACCEPTED MANUSCRIPT Table 5: Calculated equilibrium r-w and fsp-w, measured r-w and fsp-w, and W/R ratios for well 73-19. W.R. = whole-rock; Eq = equilibrium; FSP = feldspar. W.R. 18O

Eq 1 r-w

Measured 2 r-w

FSP 18O

1 fsp-w

Measured 2 fsp-w

996 886 877 822 722 490 469 207 195 42 -49 -79 -119 -143 -213 -313 -395 -440 -544 -565 -586

6.98 6.61 6.37 7.48 5.52 7.57 6.36 7.16 7.41 6.07 6.63 5.91 4.61 2.38 6.3 5.94 6.85 5.13 5.14 3.3 3.3

16.10 14.08 14.03 13.35 11.78 8.85 8.62 6.37 6.25 5.47 5.12 5.12 4.92 5.00 4.69 4.55 4.39 4.26 4.63 4.95 4.86

12.54 12.17 11.93 13.04 11.08 13.13 11.92 12.72 12.97 11.63 12.19 11.47 10.17 7.94 11.86 11.50 12.41 10.69 10.70 8.86 8.86

7.79

17.43

13.35

8.12

14.47

13.68

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6.14

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Eq

5.42

11.70

5.18

12.82

4.69 4.69

11.76 11.90

6.20 6.34

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7.26

W/R closed

W/R Open

0.08 0.13 0.16 0.01 0.28 0.00 0.16 0.06 0.02 0.20 0.13 0.22 0.40 0.70 0.17 0.22 0.10 0.33 0.33 0.57 0.57

0.08 0.12 0.15 0.01 0.24 0.00 0.15 0.05 0.02 0.18 0.12 0.20 0.33 0.53 0.16 0.20 0.09 0.28 0.28 0.45 0.45

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73-19 ID

Equilibrium r-w and fsp-w fractionation values calculated at pre-production reservoir temperatures using the experimental fractionation factor for An50-water and An25-water (O’Neil and Taylor, 1967), respectively. 2 Measured wr-w and fsp-w are the differences between the measured 18O values of whole-rock or feldspar, and the measured 18O value of the current geothermal fluid (-5.56 ‰). 1

consistent with fluid flow that is focused along discrete, higher permeability fracture zones. In all three wells, the elevations of current reservoir fluid production, which are marked by zones of lost circulation and high permeability, correspond to one of these intervals of significant 18O/16O depletion—and high W/R ratio (Fig. 7). Person et al. (2006) inferred the basic structure of the regional groundwater system that could have produced the Coso geothermal system. However, the results of our study show that the 18O/16O depletions recorded in the host rocks of the West Flank are poorly correlated with 32

ACCEPTED MANUSCRIPT pre-production temperature patterns and better correlated with known intervals of high permeability (e.g. current fluid production zones)(Fig. 6). Further, geochemical and enthalpy

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data from well tests suggest that the Coso field is divided into a series of hydrologically isolated

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(fault bounded?), single-phase reservoirs (Williams and McKibben, 1990). These relationships and results suggest that significant permeability structure has been produced within the West

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flank reservoir by an array of permeable and impermeable fault/fracture zones. As a result, the

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flow patterns responsible for the localized 18O/16O depletions measured in the West Flank host rocks are likely fracture-controlled and channelized.

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Preferential fluid flow within higher permeability fracture zones, with accompanying increase in fluid-rock interaction and 18O/16O depletion, is likely a common feature of active

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hydrothermal systems developed in crystalline rocks. For example, Smith and Suemnicht (1991)

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measured significantly lower δ18O values (< -1‰) in what are described as fractured tuffs (well

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RDO-8) compared to values in less fractured rhyolitic flows and tuffs in other wells at comparable depth in the Long Valley system, consistent with fracture-induced increases in permeability, fluid-rock interaction, and 18O/16O depletion within the fractured tuffs. Two clear examples of how fractures and faults define zones of high permeability and control patterns of fluid flow and the extent of 18O/16O depletion in fossil hydrothermal systems are the study of the Eocene Sawtooth Ring Zone within the Idaho batholith by Criss and Taylor (1983) and the wellexposed Lake City caldera system, CO by Larson and Taylor (1986b).

4.3. The Extent of Isotope Exchange Equilibrium Accompanying Fluid-Rock Interaction Because the pre-production temperatures of the reservoir fluids are known, the equilibrium oxygen isotope fractionations between feldspar and water (equilibrium fsp-w) can be

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equilibrium between the geothermal fluid and feldspar in the reservoir host rocks. If feldspar has

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incompletely exchanged oxygen isotopes with the low 18O geothermal fluid at temperatures

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below 500 oC, measured fsp-w values will be greater than the equilibrium fsp-w values. The greater this difference, the farther from exchange equilibrium is the feldspar. The equilibrium

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fractionation factors are calculated at temperatures measured in the wells prior to production, using the experimental calibration for oxygen isotope fractionation between plagioclase and

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water from O’Neil and Taylor (1967) and a plagioclase composition of An25 (average of measured feldspar compositions in the reservoir rocks; Etzel, 2016).

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Figure 9a depicts the measured 18O values of feldspar plotted against the difference between measured fsp-w and equilibrium fsp-w values. This figure illustrates that the measured

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fsp-w values for most of the sampled intervals are larger than equilibrium fsp-w values at measured (pre-production) temperatures and plot to the right of the dashed line (where measured fsp-w = equilibrium fsp-w) in figure 9a. This difference indicates that feldspars from most of the sampled intervals in these three wells have not completely exchanged with the geothermal fluid. Two groups of samples with measured fsp-w values less than equilibrium fsp-w are exceptions. The three feldspar samples with 18O values >7.0 ‰ come from the shallowest, lowest temperature intervals of 33A-7 and 73-19 and have experienced little or no 18O/16O depletion relative to their primary 18O values. As a result, these feldspars have not experienced any significant interaction with geothermal fluid, and are far from exchange equilibrium with the geothermal fluid (large measured fsp-w). However given the low temperatures for these three samples, equilibrium fsp-w values are even larger. The second exception where measured fsp-w 34

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Figure 9: A) The difference between measured fsp-w and equilibrium fsp-w plotted against the measured 18O value of each feldspar separate. B) The difference between measured fsp-w and equilibrium fsp-w plotted against measured temperature (pre-production) for the feldspar sample. Equilibrium fsp-w is calculated at pre-production temperatures. The dashed line represents equilibrium (measured fsp-w = equilibrium fsp-w) between analyzed feldspar and the preproduction reservoir fluid. Feldspars plotting to right of the dashed line have incompletely exchanged oxygen isotopes with the reservoir fluid.

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values are less than equilibrium fsp-w values are the most 18O/16O depleted feldspars from 33A-7 and 68-6. These two samples are discussed in the following section. Figure 9a also shows that in general, the lower the 18O value of the feldspar, the closer the feldspar approaches exchange equilibrium with the geothermal reservoir fluid (e.g., measured fsp-w approaches equilibrium fsp-w). This positive correlation indicates that the feldspar moves progressively toward oxygen isotope exchange equilibrium with the reservoir fluid as the extent of fluid-rock interaction increases (as indicated by progressively greater 18O/16O depletion in the feldspar and higher calculated W/R ratio). Experimental studies and theoretical considerations indicate that the rate of oxygen isotope exchange accompanying the types of surface reactions responsible for producing the hydrothermal alteration minerals observed at Coso Hot Springs increases with increasing surface area (Cole and Chakraborty, 2001). The positive correlation

35

ACCEPTED MANUSCRIPT between the progressive 18O/16O depletion in the feldspar and the increasing approach to oxygen isotopic exchange equilibrium (measured fsp-w approaches equilibrium fsp-w) observed for most

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of the feldspar samples suggests that increases in permeability of the reservoir host rock produce

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an increase in surface area of the host rock exposed to geothermal fluid, promoting fluid-rock

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interaction and isotopic exchange.

Rates of isotopic exchange accompanying hydrothermal alteration also increase

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dramatically with increasing temperature (Cole and Chakraborty, 2001). However, there is no systematic correlation of temperature with the extent of exchange equilibrium between the

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feldspar and reservoir fluid within individual wells (Fig. 9b). For feldspar samples from 68-6, the excess of measured fsp-w compared to equilibrium fsp-w value is clearly independent of

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temperature from 180oC to 265oC; for example, across the temperature range of 225-265 oC the difference between the measured fsp-w and equilibrium fsp-w decreases from 6.3 to -2.0 and then

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increases to 5.9. A similar situation holds for samples from 33A-7 over the temperature interval from 190 oC to 250 oC; the difference between measured and equilibrium fsp-w values changes from 3.6 down to -1.7 and then abruptly increases to 5. 4.4 Thermal and oxygen isotope evolution of geothermal fluids in the Coso system Measured 18O values of the feldspars are plotted as a function of the pre-production temperature (as measured in the wells) in figure 10. Curved lines define the calculated 18O values of feldspar in equilibrium with a range of 18O values of water, including the preproduction reservoir fluids (solid line). It is apparent that the analyzed feldspars from higher temperatures (>200oC) in 73-19 plot well above the equilibrium curve for the reservoir fluid, and are far from exchange equilibrium with the reservoir fluid. These feldspars have experienced limited oxygen isotope exchange, either from kinetic barriers to isotopic exchange (from lower

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Figure 10: Measured 18O values of feldspars plotted as a function of pre-production temperatures (as measured in the wells). Curved lines define the calculated 18O values of feldspar in exchange equilibrium with a range of 18O values of water (values specified next to each curve), including the pre-production reservoir fluids in these wells (solid line), as a function of temperature. Arrows indicate changes in temperature or 18O of reservoir fluid needed for the most 18O-depleted feldspar in 68-6 and 33A-7 to achieve oxygen isotope exchange equilibrium with the geothermal reservoir fluids.

temperatures in the past?) or from limited physical contact between the reservoir fluid and rock owing to lower permeability, or both.

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ACCEPTED MANUSCRIPT 38 Many of the feldspar samples from 68-6 and 33A-7 have experienced greater 18O/16O depletion and therefore more extensive oxygen isotope exchange, but have not equilibrated with

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the reservoir fluid (Fig. 10). In contrast, the two feldspars from 33A-7 and 68-6 with the lowest

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18O values plot below the equilibrium curve for the reservoir fluid. For these two feldspar

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samples, measured fsp-w values are less than equilibrium fsp-w values by -1.66 (33A-7_-1138 m) and -1.99 (68-6_-1552) (Fig. 9a). At pre-production temperatures, incomplete isotopic exchange

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would produce measured fsp-w values greater than the equilibrium Δfsp-w values. Therefore, these

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discrepancies cannot result from incomplete oxygen isotope exchange between the geothermal fluid and the reservoir rocks, because of either slow isotopic exchange kinetics or lack of

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physical contact between fluid and feldspar crystals. Measured fsp-w less than equilibrium fsp-w

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requires that the isotopic exchange took place either at higher temperatures and/or with a reservoir fluid of lower 18O value. Figure 10 shows that temperatures would need to increase by

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55 oC (to 307 oC) and 42 oC (to 288 oC) in 68-6 and 33A-7, respectively, for the most 18O/16O– depleted feldspars from each well to be in exchange equilibrium with the reservoir fluid. Alternatively, the 18O values of geothermal fluids would need to decrease by 1.98 ‰ and 1.66 ‰ in 68-6 and 33A-7, respectively (Tables 3,4; Fig. 10). A lower δ18O value of the reservoir fluid in the past implies a decrease in the initial δ18O value of the meteoric water source, and/or an increase in the W/R ratio (and permeability) for these sections of geothermal system. Oxygen isotope disequilibrium between 18O/16O depleted rocks and current reservoir fluids at present temperatures has been noted in other active hydrothermal systems as well. Sturchio et al. (1990) documented oxygen isotope disequilibrium between many of their samples of silica minerals from fractures and vugs in drill core from Yellowstone with present-day thermal waters at current temperatures. They attributed this disequilibrium to 18O/16O enrichment

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ACCEPTED MANUSCRIPT 39 of thermal fluids from the transient effects of decompressional boiling and/or increased fluidrock interaction/exchange in newly-formed fractures. Smith and Suemnicht (1991) attributed

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discrepancies between measured and calculated 18O values of host rocks in parts of the Long

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Valley system to record fluid-rock interaction (and oxygen isotope exchange) at higher

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temperatures in the past. These discrepancies noted for Yellowstone, Long Valley and Coso reflect the transient nature of the thermal and fluid-rock interaction regimes in active

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hydrothermal systems.

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5. CONCLUSIONS The patterns and extent of 18O/16O depletions measured in the reservoir rocks of the Coso

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Hot Springs geothermal system, combined with clay mineral alteration patterns, record

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significantly varying degrees of fluid-rock interactions. These patterns provide insights into the

following: 

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role of fracture-controlled permeability in this geothermal system. Our results indicate the

Clay mineral distribution in three wells along the West Flank is not always closely correlated with the distribution expected from preproduction temperatures measured in these wells. There is, however, a general correlation between greater and more widespread 18O/16O depletion in the northern part of the West Flank with a thinner smectite zone and disappearance of the illite/smectite zone. This correlation suggests that clay mineral distributions in the Coso Hot Springs field is a function of both the extent of fluid-rock interaction and temperature.



Whole-rock and feldspar 18O analyses define patterns of 18O/16O depletion in the West Flank distinct from those expected from regional metamorphism. These oxygen isotope analyses show that the reservoir rocks have experienced a wide range of 18O/16O

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ACCEPTED MANUSCRIPT 40 depletion, from negligible depletion below primary 18O values of +7.5 ‰ to 18O values as low as -6 ‰ for feldspar. This large range indicates that the reservoir rocks have

Detailed δ18O measurements of whole rock and feldspar samples from three wells along

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

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experienced significant variation in the extent of fluid-rock interaction.

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the West Flank identify a limited number of localized intervals of extensive 18O/16O depletion where the host rocks have interacted and exchanged isotopically with

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significant quantities of geothermal fluid (high calculated W/R ratios). The local zones of maximum 18O/16O depletion in each well correspond closely with the depths of the

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current production zones (intervals of high permeability). Rocks between these intervals have experienced much less exchange. These patterns likely reflect that permeability in



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the reservoir host rock is fracture-controlled. Most analyzed feldspars have 18O values too high to have completely exchanged with

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the reservoir fluid. The analyses also show that in general, the lower the 18O value of the feldspar (the greater the extent of fluid-rock interaction), the closer the feldspar approaches isotope exchange equilibrium with the geothermal fluid. This positive correlation suggests that fracture-induced increases in permeability of the reservoir host rock increase mineral/grain surface area exposed to geothermal fluids, promoting fluidrock interaction and isotopic exchange. 

There is no systematic correlation of temperature with either the extent of 18O/16O depletion or the extent of exchange equilibrium between the feldspar and reservoir fluid in the West Flank. This is particularly the case for 73-19, where rock host has experienced only moderate and highly localized 18O/16O depletion despite having the highest measured temperatures in the three wells. Thus in the West Flank at Coso Hot

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ACCEPTED MANUSCRIPT 41 Springs (a crystalline-rock-hosted geothermal system), permeability is as, or more important than temperature in driving fluid-rock interaction and isotopic exchange. The three feldspar samples with the lowest 18O values are all from current production

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

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zones in the reservoir. These feldspars are highly depleted in 18O/16O and therefore have experienced extensive fluid-rock interaction and isotopic exchange. However, two of

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these feldspar separates have measured fsp-w < equilibrium fsp-w which requires that

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they have equilibrated with the pre-production reservoir fluid at higher temperature, or

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with a lower 18O reservoir fluid. These discrepancies suggest that the West Flank of the Coso geothermal system was hotter and/or characterized by higher permeability (higher

The results of this study corroborate those of a number of previous oxygen isotope

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W/R ratio) during fluid-rock exchange in the past.

studies of (mostly) fossil hydrothermal systems, in documenting the connection between

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the extent of 18O/16O depletion, the extent of fluid-rock interaction and permeability in crystalline rock hosts to hydrothermal systems. These connections suggest that mapping patterns of 18O/16O depletion can be another exploration tool to define permeability structure, identify higher permeability zones, and constrain fluid flow patterns – past and present – in active hydrothermal systems, and to help facilitate development of enhanced geothermal systems.

6. ACKNOWLEDGEMENTS A U.S. Department of Energy (DOE) grant to Moore is gratefully acknowledged. TerraGen has provided access to well cuttings, and a variety of geological, geochemical and geophysical data on the Coso Hot Springs geothermal system. We thank Shaul Hurwitz and one

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ACCEPTED MANUSCRIPT 42 other anonymous reviewer for their thoughtful and constructive comments that helped enhance

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the overall quality of this paper. We thank Clay Jones for his assistance with XRD analyses.

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Sturchio, N.C., Keith, T.E.C., and Muehlenbachs, K. (1990) Oxygen and carbon isotope ratios of hydrothermal minerals from Yellowstone drill cores: Journal of Volcanology and Geothermal Research, v. 40, p. 23-37 Taylor, H.P., Jr. (1971) Oxygen isotope evidence for large-scale interaction between meteoric ground waters and Tertiary granodiorite intrusion, western Cascade Range, Oregon: Journal of Geophysical Research, v. 76, p. 7855-7874 Taylor, H.P., Jr. (1973) 18O/16O evidence for meteoric-hydrothermal alteration and ore deposition in the Tonopah, Comstock Lode, and Goldfield Mining districts Nevada: Economic Geology, v. 68, p. 747-764 Taylor, H.P., Jr. (1974) The application of oxygen and hydrogen isotopes studies to problems of hydrothermal alteration and ore deposition: Economic Geology, v. 69, p. 843-883 Taylor, H. P., Jr. and Forester, R.W. (1979) An oxygen and hydrogen isotope study of the Skaergaard intrusion and its country rocks: A description of a 55-M.Y. old fossil hydrothermal system: Journal of Petrology, v. 20, p. 355-419 Truesdell, A.H. (1984) Stable isotopes in hydrothermal systems in Fluid-mineral Equilibria in Hydrothermal Systems, J.M Robertson (ed): Reviews in Economic Geology, v. 1, p. 129-141 Valley, J.W. and Graham, C.M.,1996, Ion microprobe analysis of oxygen isotope ratios in quartz from Skye granite: healed micro-cracks, fluid flow, and hydro-thermal exchange. Contributions Mineralogy Petrology, vol. 124, 3/4, p. 225-234

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ACCEPTED MANUSCRIPT 47 Valley, J.W. and Kita, N.T., 2009, In situ Oxygen Isotope Geochemistry by Ion Microprobe, In: Fayek, M. (ed) MAC Short Course: Secondary Ion Mass Spectrometry in the Earth Sciences, vol. 41, p. 19-63

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Valley, J. W., Kitchen, N., Kohn, M. J., Niendorf, C. R., and Spicuzza, M. J. (1995) UWG-2, A garnet standard for oxygen isotope ratio-Strategies for high precision and accuracy with laser heating: Geochimica et Cosmochimica Acta, v. 59, p. 5223-5231

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White, A.F. and Peterson, M.L. (1991) Chemical equilibrium and mass balance relationships associated with the Long Valley hydrothermal system, California, U.S.A.: Journal of Volcanology and Geothermal Research, v. 48, p. 283-302

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Williams, A. E. and McKibben, M. A. (1990) Isotopic and chemical constraints on reservoir fluids from the Coso geothermal field, California: Proceedings, 14th Workshop on Geothermal Reservoir Engineering, Stanford, CA, p. 1545-1552

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Wilson, C.K., Jones, C.H., and Gilbert H.J., (2003) A single-chamber silicic magma system inferred from shear-wave discontinuities of the crust and uppermost mantle, Coso geothermal area, California, Journal of Geophysical Research, v. 108, B5

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APPENDIX A

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PREEXISTING WHOLE-ROCK OXYGEN ISOTOPE DATA AND UWG-2 DATA

513 preexisting whole-rock oxygen isotope analyses from 52 wells throughout the Coso

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system, made available by Terra-Gen, have been compiled into Table A.1. Every measurement

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was made at Southern Methodist University. These data helped determine the well to study in detail for this project (68-6, 33A-7 and 73-19).

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Table A.1: 513 Preexisting whole-rock oxygen isotope analyses from the Coso system. Elevation (in meters) is relative to mean sea level; 18O values are in ‰ notation, relative to VSMOW. 18O (‰) 8.03 7.21 4.23 1.22 7.54 6.07 8.47 7.38 9.32 6.74 6.82 7.15 7.52 6.04 7.88 7.70 7.66 6.92 7.17 7.02 7.98 8.19 7.50

Elevation

 O (‰)

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Elevation 1162 1089 940 662 498 402 254 117 102 -62 -167 -321 -406 -521 -559 -650 -751 -868 -939 -995 1079 1018 866

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Well BLM 84-30

BLM33B-19

NVY51A-16

Well Well

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744 604 442 373 225 70 -75 -214 -291 -436 -510 -587 -759 -840 -925 -1062 -1189 -1292 -1353 985 875 687 482 324 Elevation 223

7.50 7.28 8.43 6.83 7.79 8.55 7.65 6.51 8.06 5.12 7.03 5.85 7.43 7.55 5.54 5.79 6.77 5.10 4.48 6.91 9.60 6.84 9.60 7.66 18O (‰) 6.66

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Elevation 855

736 669 556 477 321 204 105 31 -61 -165 -230 -365 -454 -582 -707 -760 -830 -903 -1020 -1091 -1188 -1265 -1327 -1407 -1463 -1552 -1593 -1670 1091 978 858 768 734 659 1095 970 848 714 594 492 375 252 88 -34 -169 -279 -429 -512

3.32 6.35 6.60 5.16 4.26 5.33 5.31 4.77 4.45 2.72 3.77 4.39 5.22 5.15 3.82 4.59 3.56 2.99 1.48 0.99 0.57 0.35 0.66 4.02 -0.52 -4.06 0.52 0.57 6.20 6.72 6.83 6.23 5.72 5.48 7.78 7.82 7.41 7.17 7.55 6.78 7.43 7.23 7.11 6.14 7.61 5.88 6.56 5.96

Elevation

18O (‰)

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7.44 7.24 5.94 5.00 7.63 5.83 6.60 9.44 6.60 7.23 6.33 5.71 4.51 6.65 4.41 2.70 9.07 7.60 7.38 6.63 6.08 4.95 6.18 7.81 5.89 5.90 7.81 6.39 8.91 6.97 7.82 6.86 8.82 7.35 4.55 7.41 6.61 6.11 5.87 7.10 8.60 7.69 5.38 6.77 4.77 5.91 4.36

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NVY34A-9

90 -85 -88 -250 -347 -430 -546 -604 -692 -765 -889 -926 -1044 -1167 -1225 -1297 1041 946 779 645 526 350 293 123 93 -7 -63 -133 -180 -238 -324 -412 -494 -578 -664 -806 -931 -1002 -1075 -1166 -1308 -1416 -1477 -1558 -1587 -1631 1022

NVY73A-7

BLM 58A-18

18O (‰) 4.03

Well

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BLM 66-6

BLM 88-1

BLM 88-1RD

Well

Elevation 284

121 -61 782 621 392 69 -163 -406 -542 -827 -1173 -1362 963 734 543 394 201 61 -195 -343 -562 -866 1006 789 596 378 107 -94 -324 -482 -697 1172 1041 870 719 584 471 323 182 830 501 230 19 -255 -586 -816 -1024 -1234 -1397 Elevation 976

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BLM 23A-19

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NVY 63A-18

NVY 41B-8

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NVY 76B-18

7.02 5.47 6.26 5.06 3.47 6.15 7.43 6.03 6.16 6.94 5.9 5.8 4.73 5.06 4.74 7.34 7.89 7.48 7.08 4.41 4.52 7.42 5.99 7.08 9.8 8.54 1.62 4.48 4.12 4.96 5.02 7.08 6.78 4.21 4.38 3.02 3.62 4.28 0.37 2.83 0.23 7.67 6.79 7.24 7.95 5.41 7.50

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BLM 43-7

-620 -701 -792 -883 -983 -1080 -1228 -1291 938 825 752 664 569 499 426 1076 972 865 680 375 132 -103 -217 -417 -529 -730 -828 -897 -1019 -1153 -1189 893 694 529 398 239 97 -48 -368 -459 -583 1071 913 821 663 553 399

NVY 66-7

NVY 63B-18

NVY 63B-18D

18O (‰) 5.55

Well NVY 78B-6RD

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6.67 1.75 7.00 8.35 9.00 6.82 7.60 6.55 4.50 6.09 5.68 5.32 6.88 6.18 5.54 4.85 5.17 5.74 6.22 4.39 4.39 4.47 7.06 6.54 6.58 7.15 6.65 6.65 7.56 9.47 7.66 6.07 7.88 6.24 6.29 4.78 5.62 5.60 6.05 6.50 5.89 6.06 2.42 5.13 4.83 6.13 6.97 4.70 4.09 18O (‰) 7.66

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NVY 13A-16

NVY 64-16 NVY 64-16RD

Well

Elevation

-1526 697 557 345 220 75 -297 -421 -667 -912 -1122 935 630 338 22 -282 -585 -888 -1155 -1466 -1741 -433 -732 -1052 -1271 -1492 884 776 566 484 352 189 -2 998 668 552 396 177 55 -95 -223 -362 379 94 -242 -462 -724

5.25 5.42 5.61 4.28 5.79 5.54 4.52 7.32 5.77 3.29 2.82 4.52 7.07 6.92 7.11 7.21 6.83 6.51 4.65 5.48 5.78 6.35 5.28 4.23 4.60 5.81 5.55 4.44 6.98 6.17 6.98 7.39 2.75 7.40 5.29 6.14 4.84 5.00 4.65 6.31 5.85 5.75 6.98 6.77 7.90 7.12 5.17

Elevation -942

18O (‰) 7.14

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NVY 38-9

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NVY 23A-17

BLM 54-7RD

NVY 38A-9

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BLM CGEH1

6.97 5.08 3.40 3.26 3.25 4.06 3.23 2.18 6.02 6.36 7.88 6.26 6.87 7.02 7.45 6.46 5.62 5.59 6.05 6.30 7.20 5.90 6.82 5.73 6.18 4.63 4.28 8.45 7.30 7.29 7.84 7.11 7.28 7.04 7.04 7.37 7.01 6.41 7.38 6.24 7.05 7.78 8.05 7.63 7.58 5.84 3.22 2.05

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NVY 78B-6ST

706 485 243 47 -170 -347 -615 -806 1176 1063 908 756 633 436 285 84 -101 1023 767 648 453 243 -70 -343 -583 -814 -1066 927 713 595 377 262 121 -104 -303 -538 -769 888 690 258 186 -14 -298 -536 -708 -928 -1185 -1298

NVY 47A-8RD

NVY 78-7

BLM 46A-19RD

Well

18O (‰)

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Navy I 41A-8

Navy II 81A-18 Navy II 81A18RD

Navy II 67C-17

Well

0 -1118 823 452 171 -84 -370 -694 -1007 -1333 -1497 -1733 900 593 270 -36 -353 -658 -898 -1196 -1492 -1824 845 501 211 -109 -363 -706 -965 -1224 -1501 -1916 793 507 203 -100 -347 -650 -905 -1241 -1471

7.30 7.50 7.63 7.00 6.06 5.60 6.60 7.50 7.89 7.19 7.19 8.21 8.93 6.65 6.73 6.69 5.66 5.67 7.37 8.00 6.56 7.77 6.77 6.65 7.32 6.99 6.45 7.08 7.37 6.88 7.79 7.09 6.14 5.13 5.47 5.72 7.25 6.97 8.16 7.11 7.76 6.98 6.94 6.78 6.04 3.95 7.75

Elevation -1709

18O (‰) 6.82

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Navy II 86-17

Navy II 83-16

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Navy I 24A-8

6.51 6.71 7.56 0.45 0.41 5.22 7.02 5.20 8.92 9.15 5.81 5.94 5.23 6.24 5.46 6.05 6.59 6.28 6.15 4.41 6.12 4.08 1.52 2.90 1.42 6.34 6.80 6.81 5.65 7.76 6.94 5.72 6.81 7.32 6.60 6.94 7.67

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Navy I 87A-7

-1232 -1536 -1828 -2052 -2113 -2265 -2373 -2579 1128 1031 886 696 580 425 258 1120 931 693 462 288 155 24 -150 -272 -417 1003 815 601 425 239 83 -107 -305 -534 -733 1090 959

Navy II 83B-16

832 696 489 392 280 180 924 726 538

6.33 6.13 4.90 5.33 4.28 5.71 7.84 7.95 7.35

Navy II 64A-16

Elevation 321

18O (‰) 7.79

Well

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86 -181 -520 -701 0 -927

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BLM 47B-20

BLM 16A-20

BLM 24-20

Well

Elevation

352 63 -122 -439 -570 -735 972 682 454 183 -40 -237 -471 -680 -892 -1108 962 752 472 296 72 -53 -222 -366 -748 -936 -1074 -1133 -1260 -1319 -1441 -1513

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BLM 24-20RD

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BLM 88-20

BLM 81A-19

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BLM 52-20

7.59 8.33 7.16 7.04 7.37 7.11 6.58 7.23 7.19 6.90 8.32 6.43 5.41 5.76 6.33 6.64 6.22 6.68 7.68 8.07 7.21 7.34 6.26 6.68 7.20 6.98 7.53 5.70 6.46 6.02 5.18 6.99 7.17 7.59 6.74 6.31 6.71 7.97 7.33 8.14 6.84 4.27 6.31 6.36 7.12 7.09 7.76 7.40

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Navy II 37B-17

799 537 252 47 -125 -325 -508 -693 -1009 -1381 1000 726 522 273 38 -152 -362 -576 890 654 401 201 -6 -207 -434 -616 -838 -1082 976 790 540 337 89 -31 -258 -504 -663 1056 788 599 280 41 -200 -367 -588 -908 975 656

BLM 81A-19RD

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Navy II 67-17

NVY 38B-9

18O (‰)

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7.17 6.83 6.28 5.23 4.20 4.99 6.41 8.18 6.15 6.15 6.02 7.59 7.81 7.53 7.33 6.71 7.51 7.88 7.31 7.41 5.87 7.06 6.94 6.25 5.52 5.56 4.35 6.62 5.80 5.58 5.75 6.06

ACCEPTED MANUSCRIPT All data used to construct the contour maps shown in Figure 5 are presented here in Table A.2. Measured 18O values (presented in Tables A.1 and in the main text) from

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each well were averaged over ~ 480 meter intervals to construct five contours maps.

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Standard deviations and the number of values used to calculate each averaged value are

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also reported.

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Table A.2: Averaged 18O values used for contours presented in Figure 5. The number of values per average (n) and the standard deviation are also included. Values are in per mil (‰) relative to VSOW. When n =1, the standard deviation is shown as ―n/a.‖

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Average Well 18O Range: Surface to 800 meters BLM33B-19 7.89 NVY68-6 4.20 NVY73A-7 6.58 BLM 58A-18 7.67 BLM 43-7 6.55 NVY 76B-18 7.57 BLM 66-6 7.08 BLM 23A-19 7.00 NVY 63A-18 6.88 NVY 41B-8 6.80 NVY 66-7 6.73 NVY 63B-18 6.50 NVY 78B-6RD 6.50 BLM CGEH1 6.75 NVY 23A-17 5.59 NVY 47A-8RD 5.00 NVY 78-7 7.40 Navy I 87A-7 7.96 Navy I 24A-8 6.32 Navy I 41A-8 6.57 Navy II 81A-18 7.31 Navy II 81A-18RD 6.33 BLM33B-19 6.58 Navy II 37B-17 8.32 BLM 47B-20 6.24 BLM 16A-20 7.65 BLM 24-20 7.76 BLM 81A-19 7.51 33A-7 6.35 73-19 6.86

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Standard Deviation

n

0.35 0.23 0.34 0.23 0.60 0.29 n/a n/a n/a 0.37 1.00 n/a n/a 0.99 n/a 0.78 n/a 1.87 0.38 0.33 0.52 n/a 1.06 n/a 0.31 n/a n/a n/a 1.75 0.48

3 2 3 3 3 3 1 1 1 2 3 1 1 3 1 2 1 3 2 2 2 1 4 1 2 1 1 1 5 4

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n

0.69 1.39 0.38 0.69 1.39 0.38 0.34 0.57 1.89 1.44 0.46 0.67 0.40 0.62 n/a 1.34 0.40 0.60 0.72 0.47 0.66 n/a 0.72 1.28 0.00 0.16 n/a 0.07 1.39 2.83 1.03

4 5 3 4 5 3 4 5 2 3 3 3 2 4 1 2 3 2 3 3 3 1 2 2 2 2 2 2 5 3 3

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Standard Deviation

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Average Well 18O Range: 800 meters to 320 meters BLM33B-19 7.44 NVY68-6 5.14 NVY73A-7 5.81 BLM 58A-18 7.44 BLM 43-7 5.14 NVY 76B-18 5.81 BLM 66-6 7.23 BLM 23A-19 5.25 NVY 63A-18 5.75 NVY 41B-8 5.12 NVY 66-7 8.68 NVY 63B-18 5.52 NVY 78B-6RD 6.87 BLM CGEH1 5.57 NVY 23A-17 5.89 BLM 54-7RD 6.03 NVY 47A-8RD 6.72 NVY 78-7 6.52 BLM 46A-19RD 5.10 Navy I 87A-7 6.71 Navy I 24A-8 5.42 Navy II 81A-18RD 6.98 Navy II 37B-17 5.92 BLM 47B-20 6.09 BLM 16A-20 8.14 BLM 24-20 7.29 BLM 81A-19 7.88 BLM 81A-19RD 7.36 68-6 4.88 33A-7 3.46 73-19 6.48 Range: 320 meters to -165 meters BLM33B-19 8.00 BLM 58A-18 7.02 NVY 76B-18 5.97 BLM 66-6 3.64 BLM 23A-19 7.21 NVY 63A-18 5.46 NVY 41B-8 6.65 NVY 66-7 6.05 NVY 63B-18 4.24 NVY 78B-6RD 3.33 NVY 78B-6ST 3.25 BLM CGEH1 6.51 NVY 23A-17 6.36 BLM 54-7RD 5.67 NVY 47A-8RD 5.07

0.48 0.63 2.05 0.63 0.55 0.40 0.00 n/a 2.57 0.10 n/a 0.92 0.65 0.18 3.28

3 4 2 3 2 2 2 1 2 2 1 3 2 2 2

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Average Standard Well Deviation 18O NVY 78-7 5.32 0.88 BLM 46A-19RD 6.77 n/a Navy I 87A-7 5.46 n/a Navy I 24A-8 5.22 2.09 Navy II 81A-18RD 5.00 1.01 Navy II 37B-17 6.24 0.45 BLM 52-20 6.80 0.76 BLM 47B-20 7.38 0.30 BLM 16A-20 5.56 1.82 BLM 24-20RD 6.56 0.39 BLM 81A-19RD 6.47 0.84 68-6 4.52 1.07 33A-7 3.59 0.59 73-19 5.74 1.75 Range: -165 meters to -645 meters BLM33B-19 6.51 1.12 BLM 58A-18 6.36 0.54 NVY 76B-18 7.62 1.96 BLM 66-6 1.14 1.46 BLM 23A-19 5.53 1.45 NVY 63A-18 5.00 1.06 NVY 41B-8 8.52 1.35 NVY 63B-18 4.98 0.21 NVY 78B-6ST 3.65 0.59 NVY 23A-17 5.96 0.32 BLM 54-7RD 5.92 1.98 NVY 78-7 5.80 0.07 BLM 46A-19RD 7.51 0.55 Navy I 24A-8 4.61 2.90 Navy II 67-17 6.85 0.37 Navy II 37B-17 6.45 0.33 BLM 47B-20 6.59 0.24 BLM 16A-20 6.60 0.45 BLM 81A-19RD 6.60 0.49 NVY68-6 4.63 0.69 33A-7 2.54 1.44 73-19 5.44 1.24 Range: -645 meters to -1125 meters BLM 58A-18 5.28 1.13 NVY 76B-18 4.69 2.86 BLM 23A-19 6.09 n/a NVY 63A-18 4.47 n/a NVY 41B-8 7.66 n/a NVY 63B-18D 6.55 0.59 NVY 78B-6ST 2.18 n/a NVY 23A-17 4.46 0.25

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Standard Deviation 1.59 1.39 n/a n/a n/a 1.40 1.77

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Average 18O 3.96 6.16 6.60 7.09 4.99 2.91 0.46

Well BLM 54-7RD BLM 46A-19RD Navy I 41A-8 Navy II 67C-17 BLM 24-20RD 68-6 33A-7

The 18O value of garnet standard UGWG-2 (Valley et al., 1995) was repeatedly

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analyzed during the period all samples from wells 68-6, 33A-7 and 73-19 were analyzed

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for this study. These measurements were made at the University of Wisconsin-Madison. In total, 34 measurements were made and are reported in table A.3.

5.79 18O 5.55 5.78 5.79 5.77 5.83 5.92 5.79 5.69 5.82 5.83 5.83 5.56 5.70 5.72 5.78 Avg = 5.71 2 = 0.20

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Date measured 9/4/12 9/4/12 9/4/12 9/4/12 9/4/12 9/4/12 9/4/12 9/4/12 9/5/12 9/5/12 9/5/12 9/5/12 9/5/12 9/5/12 9/6/12 9/6/12 9/6/12 9/6/12

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18O value 5.60 5.55 5.63 5.66 5.72 5.73 5.69 5.63 5.62 5.61 5.61 5.73 5.53 5.69 5.92 5.70 5.73 5.79

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Table A.3: 18O values for UWG-2. These measurements (n = 34) were made during the course of this study. All values are reported in ‰ notation relative to VSMOW.

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9/6/12 Date measured 9/6/12 9/7/12 9/7/12 9/7/12 9/7/12 4/1/14 4/1/14 4/1/14 4/2/14 4/2/14 4/2/14 4/2/14 4/3/14 4/3/14 4/3/14

ACCEPTED MANUSCRIPT APPENDEX B

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PROCEDURE AND DATA USED TO CALCULATE W/R RATIOS

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The equations and procedures for calculating water(W)/rock(R) ratios using both closed and open system box models follow Taylor (1971) and Criss and Taylor (1986). The average

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δ18O value of local meteoric water (geothermal fluid source) with similar D values (-96 to -104

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‰) to those of the pre-production reservoir fluids is ~ -13 ‰ (Williams and McKibben, 1990). The δ18O values of the pre-production reservoir fluids were determined using measured

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temperature, pressure and δ18O values of the separated vapor (steam) and liquid water sampled during pre-production flow tests in each, following the heat and mass balance calculations

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described by Truesdell (1984). Use of samples from pre-production flow tests avoids the impact

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on the calculated δ18O values of the reservoir fluids of any 18O/16O enrichment from steam loss

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during subsequent power production. In all the text, the δ18O values of the reservoir fluid refer to the calculated δ18O values of the pre-production fluids.

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ACCEPTED MANUSCRIPT Highlights

Oxygen isotope systematics in an active geothermal system

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Detailed δ18O values of whole rock and feldspar samples from 3 producing wells

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Extensive 18O/16O depletion correspond closely to current production zones

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Measured fsp-w < equilibrium fsp-w for select samples demonstrates system evolution

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