Hydrogeology of the Krafla geothermal system, northeast Iceland

Transcription

1 Geofluids (2016) 16, doi: /gfl Hydrogeology of the Krafla geothermal system, northeast Iceland E. C. POPE 1,2,D.K.BIRD 2,S.ARNOR S SON 3 AND N. GIROUD 3,4 1 Natural History Museum of Denmark, University of Copenhagen, København, Denmark; 2 Department of Geological and Environmental Sciences, Stanford University, Stanford, CA, USA; 3 Institute of Earth Sciences, University of Iceland, Reykjavık, Iceland; 4 NAGRA National Cooperative for the Disposal of Radioactive Waste, Wettingen, Switzerland ABSTRACT The Krafla geothermal system is located in Iceland s northeastern neovolcanic zone, within the Krafla central volcanic complex. Geothermal fluids are superheated steam closest to the magma heat source, two-phase at higher depths, and sub-boiling at the shallowest depths. Hydrogen isotope ratios of geothermal fluids range from 87&, equivalent to local meteoric water, to 94&. These fluids are enriched in 18 O relative to the global meteoric line by &. Calculated vapor fractions of the fluids are wt% (~0 16% by volume) in the northwestern portion of the geothermal system and increase towards the southeast, up to 5.4 wt% (~57% by volume). Hydrothermal epidote sampled from 900 to 2500 m depth has dd values from 127 to 108&, and d 18 O from 13.0 to 9.6&. Fluids in equilibrium with epidote have isotope compositions similar to those calculated for the vapor phase of two-phase aquifer fluids. We interpret the large range in dd EPIDOTE and d 18 O EPIDOTE across the system and within individual wells (up to 7& and 3.3&, respectively) to result from variable mixing of shallow sub-boiling groundwater with condensates of vapor rising from a deeper two-phase reservoir. The data suggest that meteoric waters derived from a single source in the northwest are separated into the shallow sub-boiling reservoir, and deeper two-phase reservoir. Interaction between these reservoirs occurs by channelized vertical flow of vapor along fractures, and input of magmatic volatiles further alters fluid chemistry in some wells. Isotopic compositions of hydrothermal epidote reflect local equilibrium with fluids formed by mixtures of shallow water, deep vapor condensates, and magmatic volatiles, whose ionic strength is subsequently derived from dissolution of basalt host rock. This study illustrates the benefits of combining phase segregation effects in two-phase systems during analysis of wellhead fluid data with stable isotope values of hydrous alteration minerals when evaluating the complex hydrogeology of volcano-hosted geothermal systems. Key words: basalt-hosted geothermal systems, epidote, hydrogen stable isotopes, Iceland neovolcanic zone, Krafla geothermal system, oxygen stable isotopes, water-rock interaction Received 10 July 2014; accepted 19 April 2015 Corresponding author: Emily C. Pope, Natural History Museum of Denmark, University of Copenhagen, 1350 København, Denmark. emily@snm.ku.dk. Tel: Geofluids (2016) 16, INTRODUCTION Traditionally, investigations of actively produced geothermal systems have employed oxygen and hydrogen isotope chemistry of wellhead discharge to determine the source, character and evolution of high-temperature (>200 C) geothermal fluids. Recently, Pope et al. (2014) suggested that oxygen and hydrogen isotope studies of H 2 O-bearing hydrothermal minerals are also necessary to fully characterize the fluid source and evolution in geothermal systems, and Arnorsson et al. (2007) demonstrated that phase segregation must be accounted for in data interpretation of two-phase geothermal systems. Here, we re-evaluate the hydrogeology of a volcano-hosted meteoric geothermal system in Iceland as a case study for the utility of these methods. The Krafla geothermal system is located in the volcanically active Krafla caldera, a central volcano in the northeast volcanic zone of Iceland (Fig. 1). Presently, the region hosts a 60MW e power plant (Bj ornsson 2006) and is the site of the first Iceland Deep Drilling Project (IDDP) drillhole, which encountered a shallow silicic magmatic body at 2104 m depth in 2009 (Elders et al. 2011; 2015 John Wiley & Sons Ltd

2 176 E. C. POPE et al. (A) (B) Fig. 1. Location map of the Krafla geothermal field. (A) Geothermal fields of the Krafla region (hatched areas), shown with associated central volcano, fissure swarms, and surface eruptions. G (Gaesafj oll), H (Hag ong), and K (Krafla) represent topographic highs 200 to 300 m above the elevation of the geothermal field, and potential groundwater recharge locations (G and H) or hydrologic barriers (K). Black box outlines the central geothermal field in Krafla, detailed in (B). N represents the Namafjall geothermal region. Adapted from Gudmundsson & Arnorsson (2002). (B) The Leirbotnar and Sudurhlıdar well fields in the central geothermal system at Krafla. A A denotes the cross section given in Figs 2 and 5. B B denotes the cross section given in Figs 8 and 9. Pope et al. 2013; Zierenberg et al. 2013). This is the second incident in which magma was intercepted during drilling in the geothermally active region at Krafla (Mortensen et al. 2010), and the first time it has been utilized in the production of a magma-enhanced geothermal system (Fridleifsson et al. 2015). Fluid chemistry in the Krafla geothermal system has been explored over the past thirty years in connection with geothermal energy development using fluids sampled from the discharge at wellheads (Stefansson 1981; Armannsson et al. 1987; Darling & Armannsson 1989; Arnorsson 1995). These earlier studies concluded that multiple groundwater recharge areas and the injection of magmatic volatiles into the geothermal reservoir drove variations in fluid H and O isotope compositions across the central geothermal region. However, the multiple recharge areas that would source such isotopically disparate groundwaters remain uncertain. Further, these earlier studies do not take into account isotopic fractionation effects that occur between the wellhead discharges and the aquifer fluids (i.e., fluids directly sourcing the well discharge and representing a portion of the subsurface reservoir of geothermal fluids) of excess enthalpy wells. Such fractionation results from segregation of liquid and vapor in the depressurization zone around discharging wells and partial or complete retention of the liquid phase in the subsurface (Arnorsson et al. 2007). In this study, we suggest that there is only one primary fluid source, locally derived from nearby meteoric waters from the north, consistent with the hypothesis of Arnorsson (1995). The large spatial variation in the stable isotope composition of geothermal fluids at Krafla is due to ascent of steam from a deep aquifer zone into a shallower zone of sub-boiling water and their mixing in variable proportions. In support of this hypothesis, we present new hydrogen and oxygen isotope data of geothermal aquifer fluids, calculated from the composition of wellhead discharges, and compare them to the hydrogen and oxygen isotope composition of hydrothermal epidote from drillhole cuttings (Pope et al. 2014). Epidote is a common secondary mineral of basalt-hosted hydrothermal systems, whose isotope composition is sensitive to temperature and fluid isotope composition (Bird & Spieler 2004). Such measurements provide resolution of vertical and lateral variations in the stable isotope composition of geothermal fluids that have reacted with reservoir rocks, giving spatial detail that is not possible from sampling of wellhead fluids, which provide only mean isotope values of an individual well. Using the isotope composition of hydrothermal epidote, we are able to determine the source of local geothermal

3 Krafla Hydrogeology 177 fluids and monitor their chemical evolution as they are transported through the geothermal system, and undergo mixing, boiling, phase separation, and preferential vapor transport. Combining such analyses with calculated aquifer vapor fractions and fluid isotope compositions, dissolved CO 2 concentrations, and general hydrological considerations, we present a revised hydrogeologic model for the Krafla geothermal system, which complements structural and lithological data from previous studies (e.g., Armannsson et al. 1987), and accounts for the disparate chemical features of Krafla geothermal fluids as a function of their geochemical evolution. BACKGROUND Geology of the Krafla geothermal system The Krafla geothermal system is located within the northern extension of the neovolcanic zone in Iceland, shown in Fig. 1. It is associated with the Krafla caldera, which formed within a central volcano astride a major NNE SSW trending fissure swarm ~ years ago ( Armannsson et al. 1987). Volcanism in the caldera is dominated by olivine-tholeiite basalts, but less frequent silicic eruptions (gray shaded areas, Fig. 1A) form many of the topographic highs in the region, together with subglacially erupted hyaloclastites (Sveinbj ornsdottir 1992; Jonasson 1994). Deep drillings ( 2 km) at Krafla are confined to the southeastern part of the caldera where surface manifestations of high-temperature geothermal activity (>200 C), such as fumaroles and hot altered ground, are abundant (red hatched areas in Fig. 1A). The largest and most central of these regions comprises two wellfields: Leirbotnar in the northwest and Sudurhlıdar in the southeast (Fig. 1B). A semi-linear series of explosive craters, called Hveragil, separates the fields. Approximately 2 km south of Leirbotnar is Hvitholar wellfield, which lies near the southern caldera margin. Further south, along the boundary of the central volcano another wellfield produces from the Namafjall geothermal region (Fig. 1A). These last two wellfields are minimally addressed in this study. Subsurface lithology of the Krafla geothermal system, based on analysis of drill cuttings from geothermal wells, is laterally continuous across both the Leirbotnar and Sudurhlıdar wellfields ( Armannsson et al. 1987). A generalized NW SE geologic cross section is given in Fig. 2A. Alternating layers of hyaloclastites and interglacial basalts dominate the upper m of the system. With increasing depth, basalt, dolerite, and gabbro intrusions are more pervasive, becoming the dominant lithologies below 1300 m. Minor granophyre intrusions occur throughout the system ( Armannsson et al. 1987), and silicic magma was directly encountered during drilling of wells 39 and IDDP-1 at 2571 m and 2104 m depth, respectively (Fig. 1B; Mortensen et al. 2010; Elders et al. 2011). Alteration mineralogy in the geothermal system follows temperature-dependent zonation with depth, typical for high-temperature geothermal systems worldwide. Zeolites and smectite occur in the upper few hundred meters of the system, where temperatures are <150 C. With increasing temperature and depth, alteration is progressively dominated by mixed-layered clays, chlorite, epidote, and actinolite (Fig. 2A, Armannsson et al. 1987; Sveinbj ornsdottir 1992). Measured down-hole temperatures correlate well with the stability of such index minerals, indicating they represent modern metamorphic conditions. Geochemistry of hydrothermal fluids Elemental and volatile chemistry of Krafla geothermal fluids are reviewed in detail in Armannsson et al. (1987, 1989), Arnorsson (1995), Gudmundsson & Arnorsson (2002), and Bird & Spieler (2004). Total dissolved solid concentrations of Krafla waters are ~ ppm, having ionic strengths at or below ~0.01 as computed using the WATCH program (Arnorsson et al. 1982; Bjarnason 1994). Such dilute fluids are typical for meteoric-dominated Icelandic geothermal systems, resulting from limited Cl availability in basalt-hosted systems (Arnorsson 1995). ph in aquifer fluids of individual wells is typically near neutral (~6 7), and primarily controlled by host rock dissolution and secondary mineral precipitation (Arnorsson 1995). However, shallow magmatic activity in the Krafla region has caused massive influx of magmatic gases into individual wells, resulting in ph as low as 2 in some wells ( Armannsson et al. 1987, 1989; Arnorsson 1995; Bird & Spieler 2004). Continued monitoring of gas concentrations in geothermal fluids for the last three decades has shown systematic decreases in CO 2,H 2 S and H 2 concentrations in Leirbotnar well fluids since their peak during the Krafla fires ( Armannsson et al. 1989), but an increase in CO 2 concentration of Sudurhlıdar wells, suggesting that the flux of magmatic gases from the intrusions are migrating from northwest to southeast with time (Gudmundsson & Arnorsson 2002). Oxygen and hydrogen stable isotope chemistry of fluids Stable isotope compositions of geothermal fluids in the Krafla region have been reported by Arnason (1977), Sveinbj ornsdottir et al. (1986), and Darling & Armannsson (1989). The results of these studies are illustrated in Fig. 3 by gray open symbols, shown relative to the meteoric water line (Craig 1961a), and local groundwater (dd = 87&, d 18 O = 12.3&). Early analyses from the Leirbotnar field indicated a local source for geothermal

4 178 E. C. POPE et al. Depth (m) A LEIRBOTNAR SUDURHLÍDAR A (A) Region of producing aquifers LEGEND Hyaloclastite Basaltic lavas Basalt intrusives Gabbro Felsic intrusives Zeolite-smectite zone (<200ºC) Mixed layered clay zone ( ºC) Chlorite zone ( ºC) Chlorite-epidote zone ( ºC) Epidote-actinolite zone (>280ºC) Caprock o C 400 Sudurhlídar Depth (m) Confining bed -88 (B) 340 o C 220 o C Upflow zone 300 o C 92 Upflow zone Leirbotnar Distance along profile (m) Temperature (ºC) (C) Fig. 2. Subsurface hydrogeology of the Leirbotnar and Sudurhlıdar well fields. (A) General host rock lithology (patterns) and alteration mineralogy (colors) of the geothermal system. Major subvertical faulting denoted by thick lines. The primary boundary between an upper hyaloclastite/lava lithology and a lower, less permeable intrusive lithology is also shown (blue line). This region is coincident with the depth of producing aquifers in most wells. Adapted from Armannsson et al. (1987). (B) A schematic cross section showing groundwater flow in the Krafla geothermal system, from Armannsson et al. (1987) and Darling & Armannsson (1989). Hatched lines indicate zones of decreased permeability, inhibiting fluid flow. (C) Representative temperature profiles of the Leirbotnar and Sudurhlıdar well fields based on well log data, adapted from Armannsson et al. (1987). fluids, as hydrogen and oxygen isotope compositions of the fluids were generally similar to local precipitation ( Arnason 1976, 1977; Sveinbj ornsdottir et al. 1986). Based on the similarity between d 18 O of geothermal fluids and local groundwater, but significant depletion in 18 Oof whole-rock basalts from drill-cuttings relative to MORB (d 18 O WHOLE-ROCK is 8.6 to 3.2&), Sveinbj ornsdottir et al. (1986) calculated a high water/rock ratio for the system, between ~10 and 100. Further isotope measurements by Darling & Armannsson (1989) in both Leirbotnar and Sudurhlıdar confirmed the data of Arnason (1977) for Leirbotnar. However, Darling & Armannsson (1989) determined that significant variation in the dd and d 18 O of geothermal fluids between the Leirbotnar and Sudurhlıdar fields required two distinct sources, with isotopic compositions of dd = 88&, d 18 O = 11.5&, and dd = 92&, d 18 O = 11.8&. Armannsson et al. (1987) incorporated these disparate fluid sources into the then accepted hydrogeologic interpretation of Krafla, summarized below. Previous models for fluid source and flow directions The Leirbotnar and Sudurhlıdar wellfields of the Krafla geothermal system are clearly distinguished by their subsurface temperature profiles (Fig. 2C). An inferred magma heat source ~3 7 km below the surface (Einarsson 1978) allows deep recharging groundwater to form a zone of superheated steam that is overlain by a two-phase reservoir in both wellfields, but is confined to depths below ~1000 m beneath Leirbotnar. At shallower depths geothermal fluids are sub-boiling at a near uniform temperature of

5 Krafla Hydrogeology 179 running fractures (Fig. 1A; Darling & Armannsson 1989; Arnorsson 1995). The lower dd values of the deep fluids from the Sudurhlıdar field were interpreted to indicate a different fluid source (Darling & Armannsson 1989). Groundwater recharge from the north is unlikely for the Sudurhlıdar field, as Krafla Mountain forms a hydrologic barrier between the geothermal field and the high-elevation plateau Hag ong in the northeast (Fig. 1A). With no regional highs to the south or east in the Krafla caldera, Armannsson et al. (2014) suggested Sudurhlıdar may be recharged from glacial waters formed as far as km to the south, in the Dyngjufj oll region. However, no definitive source for deep fluids in either the Leirbotnar or Sudurhlıdar wellfields has been established. Stable isotope chemistry of hydrothermal epidote Fig. 3. Hydrogen and oxygen isotope composition of geothermal fluids from Krafla. Leirbotnar is shown as circles, Sudurhlıdar as diamonds, and Hvitholar as squares. Solid symbols represent aquifer fluid compositions determined from the calculations in Tables 1 and 2. Open symbols are weighted averages of the vapor and liquid phases collected at wellheads, from Darling & Armannsson (1989) and Sveinbj ornsdottir et al. (1986). Aquifer fluids from the more southern Namafjall geothermal field are also shown as triangles (this study). Fluid data are shown relative to local groundwater and the meteoric water line (MWL) after Craig (1961a). ~205 C. Stefansson (1981) postulated that an aquiclude between upper and lower reservoirs caused the unusual temperature profile of Leirbotnar, suggesting that it pinches out eastwards beneath Hveragil. A conceptual model of the geothermal reservoir at Krafla is summarized in Fig. 2B, developed from these temperature profiles in conjunction with variations in wellhead pressure, fluid chemistry, and well-production characteristics such as transmissivity (Stefansson 1981; B odvarsson et al. 1984a,b,c; Pruess et al. 1984; Armannsson et al. 1987, 2014). Major upflow zones are identified at Hveragil and along the southeastern margin of Sudurhlıdar. The main upflow zone at Hveragil is fed by deep hightemperature fluids from the lower zone of Leirbotnar. Fluids in the upper zone of Leirbotnar are interpreted as a mixture of vapor from boiling fluids in the lower zone and local cold groundwater recharge. The Sudurhlıdar wellfield is fed by a single deep fluid source, whose primary upflow zone is in the southeast, near wells 16, 17, and 20. These fluids may mix with Leirbotnar fluids along Hveragil, the contact between the two regions. Pressure gradients across the drilled area are consistent with recharge of cold groundwater in Leirbotnar from local precipitation along Gaesafj oll, a ~ m elevation plateau northwest of the geothermal system, with flow toward Krafla concentrated along approximately N S To identify the source of geothermal fluids for the Krafla geothermal system and explain their compositional heterogeneity, we use hydrogen and oxygen stable isotope compositions of the secondary alumino-silicate mineral epidote. The presence of epidote in basalt-hosted hydrothermal systems is controlled by temperature, permeability, and fluid composition ( Arnason & Bird 1992; Arnason et al. 1993). In the Krafla geothermal system, trace amounts of epidote are found in wells at temperatures as low as 200 C, but the mineral becomes abundant, filling veins, and replacing primary plagioclase at temperatures above 260 C (Kristmannsdottir 1975, 1979). In the Leirbotnar field, the first appearance of epidote in drill cuttings is at depths of ~ m, but in Sudurhlıdar minor epidote is present at depths as shallow as 300 m (Fig. 2A, Armannsson et al. 1987). Alteration of primary minerals and glass is extensive in the porous basaltic lavas and hyaloclastites that compose the upper stratigraphy of Krafla, but in the deeper, less permeable intrusion-dominated part of the system, epidote, and other secondary minerals only form along veins and fractures (Kristmannsdottir et al. 1976). Nonetheless, its presence as a major alteration mineral over a vast spatial range in basalt-hosted geothermal systems like Krafla and its sensitivity to fluid isotope composition and temperature make epidote an invaluable tool for studying the spatial and temporal evolution of geothermal fluids (Pope et al. 2014). The temperature dependence of oxygen and hydrogen stable isotope fractionation between epidote-group minerals and water has been studied experimentally and theoretically in many studies (Graham & Sheppard 1980; Graham et al. 1980; Vennemann & O Neil 1996; Chacko et al for D/H fractionation; Matthews et al. 1983; Smyth & Clayton 1988; Smyth 1989; Zheng 1993 for 18 O/ 16 O fractionation). Here, we adopt an epidote water fractionation equation for oxygen that is a combination of experimentally determined fractionation equations for

6 180 E. C. POPE et al. zoisite quartz (Matthews et al. 1983), quartz water (Matsuhisa et al. 1979), and zoisite epidote (Kohn & Valley 1998). The resulting temperature-dependent epidote water equilibrium fractionation closely approximates experimental epidote water fractionation data of Matthews et al. (1983), illustrated in Fig. 4A as the black solid line and black filled circles, respectively. For hydrogen isotope fractionation between epidote and water, we use the equation and data presented by Chacko et al. (1999) (black line, filled circles in Fig. 4B). For discussion of the merits of applying these isotope fractionation equations, the reader is referred to Pope et al. (2009) and Pope (2011). Stable isotope analyses of hydrothermal calcite and quartz from the Krafla geothermal system by Sveinbj ornsdottir et al. (1986) show that alteration minerals are near oxygen isotope equilibrium with geothermal fluids at modern temperatures, indicating that spatial variability observed in the isotopic properties of geothermal fluids should be reflected in the isotopic composition of the alteration minerals. Thermodynamic analyses presented by Gudmundsson & Arnorsson (2005) indicate that epidote is close to chemical equilibrium with the Krafla geothermal fluids. Thus, we expect that small-scale isotopic heterogeneities of geothermal fluids will be reflected in the hydrogen and oxygen isotope composition of hydrothermal epidote with greater spatial resolution than fluids sampled at wells. Excess enthalpy well discharge Production wells in Krafla withdraw fluid from aquifers either from the shallower sub-boiling zone or from a deeper, two-phase reservoir. Fluids from the sub-boiling aquifers have liquid discharge enthalpy, that is the enthalpy of the fluids discharged at the wellhead equals that of vaporsaturated water at the aquifer temperature (Gudmundsson & Arnorsson 2002). Wells withdrawing fluid from the deeper boiling reservoir sometimes have excess discharge enthalpy, that is enthalpy of the discharged fluids is significantly higher than that of vapor-saturated water at the aquifer temperature. Excess discharge enthalpy develops when a flowing mixture of vapor and liquid segregates as it (A) (C) (B) (D) Fig. 4. Fractionation between Krafla epidote and aquifer fluids. (A) Experimental data for oxygen isotope fractionation between zoisite and water by Matthews et al. (1983) (black filled circles), and curves A (zoisite water from Matsuhisa et al. 1979; Matthews et al. 1983) and B (epidote water from Matsuhisa et al. 1979; Matthews et al. 1983; Kohn & Valley 1998). d 18 O fractionation between Krafla epidote and local aquifer fluids (Table 1) as a function of temperature shown in blue symbols. (B) Hydrogen isotope fractionation between epidote and local aquifer fluids (Tables 1 and 2) as a function of temperature, shown relative to the equilibrium fractionation curve of Chacko et al. (1999). Small black circles with error bars are experimental data by Chacko et al. (1999) used to calculate the curve. Aquifer fluid data are not available for well K-25; symbols represent fractionation between epidote and wellhead discharge fluids for this well, and only provide a minimum estimate of epidote fluid fractionation. (C and D) Oxygen and hydrogen isotope fractionation between Krafla epidote and the vapor (open symbols) and liquid components (shaded hatched symbols) of initial aquifer fluids (Table 2) as a function of temperature, relative to the same fractionation curves.

7 Krafla Hydrogeology 181 is transported through the aquifer and into wells. Some or all of the liquid water is retained in the aquifer due to its adsorption onto mineral grain surfaces by capillary forces (Gudmundsson & Arnorsson 2002; Giroud 2008; Angcoy 2010; Karingithi et al. 2010). Wet-steam wells withdrawing fluids from sub-boiling aquifers (i.e., no excess enthalpy) have approximately the same heat content and chemical and isotopic compositions in total well discharges as in the initial aquifer fluid, as there is essentially no heat exchange or reaction with the wall rock during rapid adiabatic flow to the wellhead. However, in wells producing fluids that have undergone phase segregation, neither the composition nor the heat content of the well discharge is the same as that of the initial aquifer fluid. This must be taken into account when modeling chemical and isotope compositions of the initial aquifer fluid from analysis of liquid and vapor samples collected at the wellhead. Arnorsson et al. (1990, 2007) developed an approach to model the initial vapor concentration in a geothermal aquifer before phase segregation occurs. This model is summarized briefly below and mathematically derived in Appendices A and B. Such a model is necessary to better approximate the oxygen and hydrogen isotope composition of subsurface geothermal fluids using the measured values of the wellhead discharge. METHODS Modeling aquifer fluid compositions During sampling of geothermal fluids, the vapor phase and liquid phase of the producing fluid are collected and measured for stable isotopes separately. In wet-steam wells without excess enthalpy, discharge pressure and discharge enthalpy can be measured to determine the vapor fraction of the discharging fluid (x d,v ) using steam tables (IAPWS Industrial Formulation 1997 for the Thermodynamic Properties of Water and Steam). From this, the isotope composition of the discharging fluid (equal to the aquifer fluid) can be calculated as a weighted mass balance of the vapor and liquid fractions: d d;t ¼ d d;l 1 x d;v þ d d;v x d;v ð1þ where d denotes the conventional delta notation for an isotope ratio (e.g., D/H or 18 O/ 16 O); that is, the ratio between heavy and light isotopes of an unknown sample relative to the same ratio of a standard, in this case Vienna Standard Mean Ocean Water (V-SMOW). Superscripts d,t; d,l; and d,v denote total discharge fluid, liquid fraction of the discharge fluid, and vapor fraction of the discharge fluid, respectively. In excess enthalpy wells that have undergone phase segregation, the isotope composition of the aquifer fluid (d f,t ) is the weighted mass balance of the discharge fluid (d d,t ) and the liquid water retained along the flow path from the aquifer to the wellhead (d e,l ): d f ;t M f ;t ¼ d d;t M d;t þ d e;l M e;l ð2þ where M is the mass flow rate of the aquifer (f,t), discharge (d,t), or segregated fluid (e,l). To solve this equation, it is necessary to know the temperature (T f ) and initial vapor fraction (x f,v ) of the aquifer fluid to calculate M f,t and M e,l, and the temperature and pressure at which phase segregation occur (T e and P e ) to calculate d e,l. Various chemical components within discharged aquifer fluids are used to calculate the initial aquifer vapor fraction (x f,v ) by assuming (1) specific vapor component-mineral equilibria in producing aquifers and (2) that all vapor within the producing aquifer is discharged in the wellhead fluid (Giggenbach 1980; D Amore & Celati 1983; Arnorsson et al. 1990). Here, we use the vapor component H 2, because its low solubility in liquid water makes it a sensitive monitor of vapor fraction. We consider the H 2 concentration to be controlled by local equilibrium for the mineral assemblage (Arnorsson et al. 2007, 2010): 4 3 FeS þ 2 3 Ca 2Al 2 Si 3 O 10 ðohþ 2 þ 2 3 H 2O Pyrrhotite Prehnite ¼ 2 3 Ca 2Al 2 FeSi 3 O 12 ðohþ þ 2 3 FeS 2 þh 2;aq : Epidote Pyrite ðr1þ The concentration of H 2,aq can be calculated from the temperature-dependent logk of this reaction at the aquifer temperature (T f ). T f is evaluated for each well from the average of Na/K and quartz geothermometers (Gunnarsson & Arnorsson 2000; Arnorsson & D Amore 2000), values that are supported by temperature measurements taken during thermal recovery of the wells. Once the concentration of H 2 in the aquifer fluid is known, x f,v can be determined by solving the equation (detailed in Appendix B): x f ;v ¼ m d;t s m f ;l s V 1 D f 1 f ;t s 1 ð3þ where m s denotes the concentration of vapor component s in moles/kg, V f,t = M f,t /M d,t, and Ds f is the distribution coefficient of component s in the initial aquifer. In producing aquifers, well discharges likely consist of many fluid components that have traveled different distances at different velocities from their point of origin to the well. Thus, phase segregation is likely to occur over a range of temperatures, making a precise determination of temperature and pressure at a segregation point impossible. As a first-order approximation of the system, we select a single segregation temperature (T e ) corresponding to a vapor pressure (P e ) that is half way between the wellhead pressure (P d ) and the initial aquifer pressure (P f ) to calculate d e,l.

8 182 E. C. POPE et al. Sensitivity studies (Scott 2011) indicate that modeled initial aquifer vapor fractions are correct within an order of magnitude if the pressure at which segregation occurs is approximated as a single value, unless it is close to the vapor pressure in the undisturbed aquifer. With P e and T e selected, and T f and x f,v determined from the concentration of H 2(g) in the discharged fluid, equation (2) can be solved for d f,t following the calculations detailed in Appendix A. All previous reports of the isotopic composition of geothermal fluids in Krafla assumed liquid enthalpy wells and therefore that well discharge is representative of aquifer fluids ( Arnason 1976; Sveinbj ornsdottir et al. 1986; Darling & Armannsson 1989). In this contribution, we use unpublished isotopic data of liquid and vapor sampled from Krafla wet-steam wells by Gudmundsson & Arnorsson (2002) to model the hydrogen and oxygen isotope compositions of the aquifer fluid for excess enthalpy wells at Krafla, assuming that it is caused by phase segregation. For all calculations of isotope fractionation between liquid and vapor at T f and T e, the experimentally determined fractionation equation of Horita & Wesolowski (1994) is used. Epidote stable isotope analysis Epidote was handpicked from drill cuttings from wells K- 17, K-20, K-25, K-26, K-32, K-34, and K-39 using a binocular microscope. Drill cuttings were first separated using a Frantz Separator to remove the most magnetic fraction of material. Individual grains of epidote lacking visible inclusions, rims, or coexisting minerals were selected from the remaining material. Bulk drill cuttings were sampled from the well at approximately every 2 m during drilling so that most material is from the drilled section within 2 m above the labeled depth. However, some material may be derived from higher points in the well due to caving of the drillhole walls before casing. Epidote samples were taken from drill cuttings from depths between ~900 m and the bottom of each well, the deepest being K-39 at 2362 m depth. Systematic sampling every ~100 m was attempted for each well, but due to limitations in access to drill cuttings (obtained via the Iceland GeoSurvey), grain size (cuttings larger than ~0.5 mm were necessary to separate epidote in enough quantity), and minimal alteration in low-permeability intrusive-dominated layers, this was only possible in wells K-17, K-20, K-32, and K-34. Epidote was sampled from the other wells where large enough quantities (~ mg for hydrogen analysis and mg for oxygen isotope analysis) were accessible. Hydrogen and oxygen isotope results are reported in delta notation, relative to the standard V-SMOW, and presented in parts per thousand (&; Craig 1961b). Fractionation between epidote and water for both isotopes is presented as 1000 ln(a), where 1000 ln ðaþ ¼1000 ln ½ð1000 þ d epi Þ =ð1000 þ d H2OÞŠ d epi d H2O; (Baur et al. 1978; O Neil 1986). ð4þ Stable isotope analysis of Krafla epidotes was completed at the Stanford University Stable Isotope Biogeochemistry Laboratory. Hydrogen isotopes were measured following the methods of Sharp et al. (2001). Powdered epidote samples were combusted and reduced in a Finnigan hightemperature elemental analyzer at 1450 C. The resulting H 2 gas was introduced via constant He-streaming to a Finnigan Delta Plus XL mass spectrometer. Raw isotope data were corrected using a linear regression of NBS and laboratory standards and are correct within 3&. Duplicates were run for samples where enough material was present, and 1r (standard deviation) variation is presented with the data. The laser-fluorination method outlined by Sharp (1990) was used for oxygen isotope analysis. Samples were vaporized using a CO 2 -infrared laser in a vacuum chamber in the presence of bromine pentafluoride. Oxygen gas released from the samples was directly fed into a dual inlet Finnigan MAT 252 mass spectrometer. Initial results were corrected relative to the UWG-2 garnet standard (Valley et al. 1995) and laboratory quartz standard (L 1 = &). Offsets in the data relative to the values for these standards showed no drift over a single run and varied from 0.3 to 0.8&. After being corrected for offset, measured values of standards were within 0.1& of accepted values, less than the analytical error of 0.2&. Isotope compositions of fluids in equilibrium with epidote were determined using temperatures approximated from the temperature profile of individual or proximal wells, obtained courtesy of the Iceland GeoSurvey. dd FLUID was calculated based on the equation 1000 ln a D EPIDOTE WATER ¼ 9:3000ð106 =T 2 Þ 61:90; ð5þ from Chacko et al. (1999). d 18 O FLUID was calculated based on the following equation, 1000 ln a 18O EPIDOTE WATER ¼ 1:53ð106 =T 2 Þ 3:31; ð6þ compiled from the fractionation experiments of Matsuhisa et al. (1979), Matthews et al. (1983) and Kohn & Valley (1998) as detailed above (Section Stable isotope chemistry of hydrothermal epidote ). RESULTS Fluid chemistry Tables 1 3 present data characterizing initial aquifer composition and physical conditions within the Krafla geothermal system that are based on isotopic, gas species, and

9 Krafla Hydrogeology 183 elemental analyses of fluids discharged at the wellhead. In Table 1, dd and d 18 O of the discharged fluid and the calculated initial aquifer fluid are shown. The initial aquifer fluid is a weighted sum of the isotope compositions of the liquid and vapor fractions of the two-phase aquifer fluid reported in Table 2 before any phase separation. Wellhead fluids have isotope compositions comparable to values measured by Darling & Armannsson (1989) and Sveinbj ornsdottir et al. (1986). Aquifer fluids, however, have d 18 O values elevated relative to wellhead fluids by as much as 1.7&. This is apparent in Fig. 3, where the aquifer fluids calculated in this study (solid symbols) are compared to wellhead isotope data of Darling & Armannsson (1989) and Sveinbj ornsdottir et al. (1986). As illustrated in Fig. 3, computed initial aquifer fluid compositions (Tables 1 and 2) are typical of geothermal waters in Iceland, in that they exhibit enrichment in 18 O relative to the meteoric water line due to reactions with unaltered basaltic rocks. An observable difference between Leirbotnar and Sudurhlıdar remains in the isotope composition of initial aquifer fluids, with Sudurhlıdar fluids having lower dd values, but slightly higher d 18 O relative to Leirbotnar (Fig. 3). Additionally, samples analyzed from well K-21 in Hvitholar (squares), approximately 2 km south of Leirbotnar, and the Namafjall geothermal system (triangles), located approximately 8 km south of the central Krafla geothermal system (Fig. 1), have significantly lower dd values than Leirbotnar or Sudurhlıdar, with Namafjall also showing the most enrichment in 18 O relative to the meteoric water line (Fig. 3, Tables 1 and 2). Table 2 H and O isotope distribution between liquid and vapor phases of aquifer fluids in excess enthalpy* geothermal wells. Well No. Phases in initial fluid aquifer d 18 O v (e,v) d 18 O l (e,l) dd v (e,v) Total fluid dd l d 18 O fluid (e,l) x e,v (f,t) dd fluid (f,t) *For liquid enthalpy wells, x e,v = 0, d 18 O f,t = d 18 O d,t and dd f,t = dd d,t (reported in Table 1). x = vapor mass fraction in the aquifer before phase separation. Table 1 Hydrogen and oxygen isotope composition of aquifer fluids in Krafla geothermal wells. Raw data from discharged fluids Total discharge Initial fluid aquifer Well No. d 18 O v (d,v) d 18 O l (d,l) dd v (d,v) dd l (d,l) T d P d h d,t d 18 O fluid ( C) (bar) (kj/kg) x d,v (d,t) dd P e fluid (d,t) (bar) T f ( C) P f (bar) d 18 O fluid (f,t) dd fluid (f,t) 5* * * * * N4* N4* N N N N *Liquid enthalpy wells. Hvitholar. See Section Modeling aquifer fluid compositions for explanation of all superscripts and subscripts.

10 184 E. C. POPE et al. Table 3 Measured and calculated aquifer properties for Krafla geothermal wells. Well No. Aquifer Temp. ( C) P d * (bar) h d,t (kj/kg) h f,t (kj/kg) P e (bar) H f,t 2 (mmol/kg) Aquifer fluid vapor Individual (wt%) Average (vol%) 5** % 9** % % % % % % % % % % % , ** % % ** % % % % *Sampling pressure. Discharge enthalpy. Specific enthalpy of initial fluid. Selected segregation pressure. Volume determined from International Steam Tables (Wagner & Kretzschmar 2008). **Liquid enthalpy wells. Aquifer temperatures and vapor fractions of aquifer fluids by mass fraction (wt%) and volume % are presented in Table 3 for several Krafla wells, including those for which isotope data were collected for the calculations in Table 1. Weight percent fractions are given for each analysis when there was more than one sample for an individual well, whereas volume percent fractions are an average of all analyses for an individual well. Also included in Table 3 are sampling pressure (P d ), discharge enthalpy (h d ), H 2 concentration measured at the wellhead, and the specific enthalpy of initial aquifer fluid (h f ) at the selected segregation pressure (P e ). These factors are required to compute in situ vapor fractions of high enthalpy wells as per the methods described in Appendices A and B. In Table 4, aquifer temperatures and vapor volumes for each well from Table 3 are ordered along an approximate northwest to southeast transect (A-A in Fig. 1B) through the Leirbotnar and Sudurhlıdar fields. Wells K-15, K-32, and K-12, shaded gray in Table 4, represent the transition zone between the two fields along Hveragil. Hydrogen and oxygen isotopes of fluids discharged at the wellhead (averaged from multiple analyses by Sveinbj ornsdottir et al.

11 Krafla Hydrogeology ; Darling & Armannsson 1989 and this study) and of the initial aquifer fluid (Table 1) are also shown. Well K-32 was drilled after the wellhead samples listed in Tables 1 3 were collected, so isotope values of the geothermal fluids were approximated using averages of data from wells located within 150 m of the well. No isotope data were available for well K-34, nor are there any nearby wells for which fluid isotope analyses are available. Finally, the dissolved CO 2 concentration of the initial aquifer fluid is shown for each of these wells in mmol/kg (Gudmundsson & Arnorsson 2002 and this study). The aquifer properties presented in Table 4 are shown graphically in Fig. 5. Several characteristics distinguish the northwestern Leirbotnar and southeastern Sudurhlıdar regions of the Krafla geothermal system. Temperature, vapor fraction and CO 2 concentrations of Sudurhlıdar wells are all higher than in the Leirbotnar geothermal area. dd of the initial aquifer fluid (solid circles) is inversely correlated with these variations, with a notable decrease of about 4& in the Sudurhlıdar portion of the geothermal field. Also shown in Fig. 5 is the isotope composition of the liquid and vapor fractions of the initial aquifer fluid (Table 2) as hatched and open circles, respectively. The liquid fraction closely mimics the composition of the initial aquifer fluid in each well, as by mass this fraction makes up >94% of the total initial fluid (Table 3). Vapor fraction deviates from total fluid dd as a function of temperature; at temperatures below 230 C, vapor is depleted in deuterium relative to the initial fluid whereas above this temperature it is enriched in deuterium. At the greatest observed temperature extremes, there is an ~3 5& difference between the total initial aquifer fluid composition and that of the vapor fraction. d 18 O is the only feature of the aquifer fluids which does not show a significant change across the geothermal system in Fig. 5. Values in Sudurhlıdar are at most ~2& higher than those of Leirbotnar, but there is significant overlap in individual measurements. However, in all wells the composition of the vapor fraction of the aquifer fluid is significantly lower than the total fluid composition, by as much as 2.7&. Epidote stable isotope composition Hydrogen isotope values for epidote from seven wells in the Krafla geothermal system and oxygen isotope values for epidotes from three wells are presented in Table 5 and shown as a function of depth in Fig. 6 (Pope et al. 2014). The total range of dd in Krafla epidotes is 19&, with variation between individual geothermal wells showing a distinct north south trend (Figs 1B, 6A). Wells with the highest dd EPIDOTE ( 108&) are K-25 and K-34, the two northern-most wells. Similarly, southern-most wells K-26 Table 4 Properties of Krafla geothermal fluids. Wells ordered from NW to SE. Leirbotnar Hveragil Sudurhlidar * 32* 12* Well No. Temp. ( C) Vapor (vol%) dd Aquifer fluid d 18 O dd Wellhead fluid d 18 O CO *Shaded wells denote transition region between Leirbotnar and Sudurhlıdar. Data not available from shown well, but approximated from nearby wells (<150 m from shown well). Based on calculated initial vapor fraction and analytical data from liquid and vapor samples collected at the wellhead, in mmol/kg. Data from Gudmundsson & Arnorsson (2002), this study.

12 186 E. C. POPE et al. has dd EPIDOTE values between 127 and 124& and a negative isotope excursion at ~1300 m depth, and the northernmost well K-34, which has dd EPIDOTE from 115 to 108& and a positive isotope excursion at ~ m. Oxygen isotope values of epidotes from wells K-20, K- 26, and K-34 range from 13.0 to 9.6& (Fig. 6B). The Leirbotnar wellfield is represented by epidote sampled at two depths in well K-26, which shows a 0.6& variation ( 11.1& to 11.7&). Well K-20, located in Sudurhlıdar, has an anomalously high d 18 O EPIDOTE value at a depth of 906 m ( 9.7&), but the remainder of the values from that well are within less than 1& of one another, between 13.0 and 12.3&. Well K-34, near the upflow zone at Hveragil, has the widest range in values ( 12.3 to 9.6&) and highest d 18 O EPIDOTE value in the geothermal system. In Fig. 6B, samples are shown relative to published whole-rock data of unaltered surface volcanics (+1.8 to +5.1&, Sveinbj ornsdottir et al. 1986; Pope et al. 2013) and altered bulk rock samples from drill cuttings in several Krafla wells ( 10.5 to 3.4&, Hattori & Muehlenbachs 1982; Sveinbj ornsdottir et al. 1986). DISCUSSION Fig. 5. Aquifer fluid chemistry across Krafla geothermal system. Aquifer temperature (T f ), vapor fraction, CO 2 concentration, and dd and d 18 Oof initial aquifer fluids (closed symbols), and the liquid and vapor fractions of those fluids (hatched and open symbols, respectively) along a NW SE transect through Leirbotnar and Sudurhlıdar geothermal fields (A A in Fig. 1B). Shaded region denotes the total range in values for each variable within the geothermal field. Wells 24 and 28 have liquid enthalpy, thus are derived from sub-boiling aquifers. Based on data in Table 4; from Gudmundsson & Arnorsson (2002), Darling & Armannsson (1989), this study. and K-39 have the lowest dd EPIDOTE values ( 125 to 127&). Drill cuttings from the Leirbotnar field (round symbols in Fig. 6) show less extensive alteration, resulting in fewer available epidote samples in each well, but the epidote that was sampled from wells K-25, K-32, and K-36 exhibit limited vertical variation relative to the other wells. Well K-32, still within Leirbotnar, but located nearest to the major upflow zone at Hveragil, has intermediate hydrogen isotope values (dd EPIDOTE = 119 to 116&) relative to the northern high and southern low extremes of wells K- 25 and K-39, respectively. Wells K-17 and K-20 in the Sudurhlıdar field (diamond symbols in Fig. 6) have a similar range in dd EPIDOTE, 122 to 114& and 123 to 116&, respectively. Both wells show a low-deuterium anomaly at 1250 to 1400 m depth. Wells lying along the inferred high-permeability zone of Hveragil (square symbols in Fig. 6) include K-39, the southernmost well that The apparent complexity of fluid dynamics in the Krafla geothermal field has been previously characterized based on variations in elemental chemistry and isotope composition of wellhead fluids, regional hydrologic gradients, and the extent of interaction between geothermal fluids and their surrounding host rock across the central Krafla caldera. In this section, we add to this discussion by assessing lateral variations in aquifer fluid chemistry, and the lateral and vertical heterogeneities of fluid isotope composition that is preserved in hydrothermal epidote. Combined, these data help resolve most observed complexities in fluid composition in this system and allow us to develop a more rigorous model for the relationship between hydrologic, geologic, and structural features within the region. Geothermal fluids aquifer properties Temperature, CO 2 concentrations, and vapor mass fractions of geothermal fluids are significantly different between Leirbotnar and Sudurhlıdar. Average aquifer temperatures, or the temperature at the location of the primary aquifer within the well, are ~43 C higher across the Sudurhlıdar region ( C) than in the Leirbotnar region ( C, Table 4, Fig. 5). Based on the temperature profiles with depth of Fig. 2C, primary geothermal fluid-feed zones in wells penetrating the Leirbotnar aquifer must be within a transition zone between the upper basalt flows and hyaloclastites (constant temperatures at ~205 C) and the lower crystalline mafic intrusions

13 Krafla Hydrogeology 187 Table 5 Isotopic composition of aquifer fluids, epidotes, and fluids in equilibrium with epidote.* Aquifer fluid composition Epidote composition Fluid in equilibrium with epidote Well No. dd d 18 O Depth (m) dd d 18 O Temp. ( C) dd FLUID d 18 O FLUID K K K K K K K *Tabulated presentation of data originally published in Pope et al. (2014). Data compiled from all proximal wells (Tables 1 and 2, this study) and are applied at every depth. Represents 1 standard deviation (1r) for samples where N > 1. Temperature at the depth of epidote formation, as determined from boiling point curve with depth (Sudurhlıdar), or temperature profile information by Iceland GeoSurvey (Leirbotnar). Aquifer fluid data not available for this well. Italicized values are approximated from wellhead fluids collected in nearby wells by Darling & Armannsson (1989) and provide only a minimum estimate. (<300 C at uppermost boundary), at about m depth. The temperature range of aquifer fluids in Sudurhlıdar would indicate that the primary aquifer feed zone is from about 700 to 1000 m depth (Fig. 2C), again just above the lithologic transition from upper basalts and hyaloclastites, and lower intrusives (Fig. 2). Gudmundsson & Arnorsson (2002) rigorously analyzed aquifer temperatures and the depth level of producing aquifers in the Krafla geothermal system. Their results are consistent with the above approximations; in both Leirbotnar and Sudurhlıdar, most wells encounter producing aquifers at or above the lithologic boundary marking the transition to crystalline intrusions: m and m, respectively. Wells K-12, K-15, and K-32, in the Hveragil region, have the highest aquifer temperatures ( C), consistent with its having enough vertical permeability that deeper feed zones also contribute to aquifer fluids. CO 2 concentrations of geothermal fluids are also 2 10 times greater in Sudurhlıdar than in Leirbotnar (Fig. 5). While CO 2 concentrations of wellhead fluids have been generally decreasing since the Krafla fires in ( Armannsson et al. 1989), escalated input of volcanically sourced CO 2 may still occur locally in the geothermal system, as there is evidence for an underlying magma body beneath the Krafla caldera. Armannsson et al. (1989) proposed that magmatic gases likely follow the vertical fractures controlling upflow of deep geothermal fluids, which are particularly concentrated near the Hveragil region. If this is the case, the CO 2 concentrations presented in Fig. 5 suggest that geothermal fluids containing a magmatic component are flowing in an approximately southeastward

14 188 E. C. POPE et al. (A) (B) Fig. 6. Hydrogen and oxygen isotope composition of hydrothermal epidote. Shown as a function of depth. Circles = Leirbotnar, diamonds = Sudurhlıdar, squares = Hveragil. (A) Hydrogen isotope compositions from seven wells. Error bars represent 1r deviation for duplicate samples. (B) Oxygen isotope compositions from three wells. Shown relative to whole-rock isotope data from unaltered surface eruptions (filled diamonds Sveinbj ornsdottir et al. 1986; Pope et al. 2013) and from altered subsurface basalts collected from drill cuttings (open diamonds Sveinbj ornsdottir et al. 1986; hatched diamonds Hattori & Muehlenbachs 1982). direction from Hveragil toward Sudurhlıdar. This flow direction is also supported by the vapor fraction data in Fig. 5. As a zone of high permeability and temperature gradient relative to the rest of the Krafla system, it is likely that significant boiling of deep aquifer fluids occurs along the vertical fractures of Hveragil. The less dense vapor phase, including magmatically derived gases, rises toward the surface along these fractures, mixing with the shallower groundwater of the aquifer feed-zones along the boundary between the upper lavas and lower intrusives, and increasing fluid vapor and dissolved CO 2 concentrations. Geothermal fluids stable isotope data With a significant influx of magmatic gases into the Sudurhlıdar system as suggested by fluid CO 2 and vapor fractions, an associated increase in dd and d 18 O of Sudurhlıdar fluids relative to Leirbotnar is expected, as magmatic H 2 O has isotope compositions of about 80 to 40& and +5.5 to +9.5&, respectively (Sheppard 1977; Taylor 1986). Instead, there is an ~4& decrease in hydrogen isotope compositions of aquifer fluids between Leirbotnar and Sudurhlıdar fields, and similar average values but increased variability in d 18 O across the Hveragil boundary (Fig. 5). Thus, isotopic variation between the Leirbotnar and Sudurhlıdar geothermal reservoirs is not controlled solely by magmatic volatile input in the southeast. The heterogeneous hydrogen and oxygen isotope compositions of the geothermal fluids may instead be due to (i) water rock interaction, (ii) mixing of multiple sources, (iii) liquid vapor fractionation and separation, or some combination of all of these. Geothermal fluids can be strongly influenced by oxygen isotope exchange with host rock (Criss & Taylor 1986), although earlier studies rejected this as a significant influence on d 18 O of Krafla fluids, given the inferred high water rock ratio of the system (Sveinbj ornsdottir et al. 1986). Their interpretation was based on wellhead fluid data that indicated d 18 O FLUID was similar to the meteoric water line (open symbols in Fig. 3). However, the recalculated isotope compositions in Tables 1 and 2 indicate that aquifer fluids have indeed been enriched in 18 O relative to meteoric water, to approximately equal values in Sudurhlıdar as in Leirbotnar (Fig. 3). We suggest that the increase of d 18 O of Krafla geothermal fluids relative to their meteoric source is due in both regions to the compounded effects of magmatic gas input and water rock interaction. The relative contribution of the other two controls mixing and boiling is better explored in the context of the epidote isotope data, which can trace both lateral and vertical changes in the fluid isotope chemistry within basalt-hosted geothermal systems. Epidote Hydrogen isotopes There is nearly 20& variation in dd of hydrothermal epidote in the Krafla geothermal system (Table 5, Fig. 6A), which reflects variations in both composition of geothermal fluids and temperature of epidote formation. In Fig. 4B, hydrogen isotope fractionation between epidote and initial aquifer fluids calculated from the discharge fluids of either the specified well or proximal wells is shown as a function of temperature, relative to the equilibrium fractionation curve of Chacko et al. (1999). Data from all seven wells are essentially within experimental error of the equilibrium fractionation curve, indicating that epidotes approach hydrogen isotope equilibrium with modern, local fluids.

15 Krafla Hydrogeology 189 However, some deviation ( 10&) exists in all wells between the measured hydrogen isotope fractionation between epidote and geothermal fluids and that expected under isotopic equilibrium conditions. These departures would suggest that the epidote is recording small-scale variations in the isotope composition of fluids over the length of the drilled well. In contrast, the dd FLUID and d 18 O FLUID used to calculate epidote fluid fractionation in Fig. 4 are averages of all measurements in Tables 1 and 2 for that well, or from wells within 150 m of those wells drilled after sampling took place (average shown in Table 5). Such values thus represent a mixture of fluid components sourced from aquifers at multiple depths. Boiling, variable phase separation and, vapor/fluid mobility could result in several per mil variations in the hydrogen isotope composition of geothermal fluids at different depths. We reject that the observed excursions of data from the Chacko et al. (1999) curve are alternatively due to (i) inaccuracy of the fractionation curve or (ii) disequilibrium between epidotes and geothermal fluids that are homogeneous in their isotopic character along the length of the well. Both of these factors would result in systematic variation from equilibrium. If the fractionation curve were inaccurate, offset would be regular in direction and magnitude. Disequilibrium would also result in a unidirectional offset, although the magnitude of offset may vary (e.g., Pope et al. 2009), and would further require that water rock ratios in the Krafla system are low enough to preclude isotopic resetting of relict alteration minerals. Unidirectional variation from equilibrium is not observed in Krafla. Instead, isotopic compositions of calcite, quartz, and whole-rock samples (Hattori & Muehlenbachs 1982; Sveinbj ornsdottir et al. 1986) indicate that water rock ratios in the Krafla geothermal system are large enough for minerals with high diffusion coefficients, such as epidote, to quickly equilibrate with modern, local fluids. Epidote Oxygen isotopes Oxygen isotope fractionation between epidotes and calculated initial aquifer fluids from wells K-20, K-26, and K-34 are tabulated in Table 5 and shown in Fig. 4A, relative to zoisite water fractionation (curve A) and epidote water fractionation (curve B). Epidote [Ca 2 (Al 3-X Fe 3+ X) Si 3 O 12 (OH)] in the Krafla system is compositionally intermediate between pure Al-endmember zoisite (X = 0), and stoichiometric epidote (X = 1). Bird & Spieler (2004) reported epidote compositions with 18 32% of pure Fe 3+ - endmember pistachite (X = ), consistent with that reported by Arnorsson et al. (2007) (X 0.8; or ~26% pistachite). Fractionation factors that are intermediate between curves A and B in Fig. 4A will thus be representative of oxygen isotope equilibrium between Krafla epidotes and water. It is apparent in Fig. 4A that geothermal epidote from Krafla is not in equilibrium with calculated initial aquifer fluid compositions. With only two exceptions, epidote aquifer fluid fractionation is >2& below equilibrium values. Rather, in Fig. 4C, oxygen isotope fractionation between hydrothermal epidote and the liquid and vapor fractions of the calculated aquifer fluid compositions for each well (Table 2) are shown relative to the fractionation curves of Fig. 4A. The same is shown for hydrogen isotopes in Fig. 4D. In both diagrams, epidote vapor fraction pairs are closer to equilibrium fractionation curves, significantly more so in regard to oxygen isotopes. We propose that the d 18 O and dd of Krafla epidote is determined by a fluid that is isotopically distinct from the average fluid sampled at the wellhead. We have calculated the oxygen and hydrogen isotope values of fluids that would be in isotopic equilibrium with individual geothermal epidotes in Table 5, using the same fractionation curves as in Fig. 4 and the measured down-hole temperatures for each well. The discrepancy between fluids in isotopic equilibrium with epidote and average aquifer fluid compositions, and the implications of these results on the source(s) and evolution of Krafla geothermal fluids are discussed in the next section. Origins of Krafla geothermal fluids Oxygen and hydrogen stable isotope compositions of fluids in equilibrium with epidote from wells K-20, K-26, and K- 34 (Table 5) are plotted in Fig. 7, where they are compared to the meteoric water line and calculated isotope compositions of the vapor phase (open red diamonds) of aquifer fluids in both Leirbotnar and Sudurhlıdar (Table 2). Most calculated values from wells K-20 and K- 34 have lower d 18 O and higher dd than meteoric waters, as do most aquifer vapor phase fluids. Natural waters with isotope compositions above the meteoric water line like these are produced almost exclusively by separation of the vapor fraction of boiled meteoric water. The excess enthalpy measured in most wells of the Krafla geothermal system indicates that phase separation and preferential mobility of the vapor phase from two-phase aquifers is common. The resulting isolation of liquid and vapor phases of the initial aquifer fluid likely occurs along natural vertical conduits in the geothermal system, such as faults and fractures, which are prevalent due to caldera formation and plate spreading. As vapor (together with magmatic gases like HCl, CO 2 and H 2 S) rises along these conduits to cooler sections of the geothermal system, it condenses to a dilute fluid or mixes with shallower, sub-boiling aquifer fluids. The condensed fluid reacts with the wallrock to form epidote by reactions with igneous plagioclase and volcanic glass. Like initial aquifer fluids, which have ionic strengths 0.01 but are saturated with respect to epidote,

16 190 E. C. POPE et al. Fig. 7. dd and d 18 O of geothermal fluids in equilibrium with epidote from wells K-20, K-26, and K-34. Shown relative to the meteoric water line (MWL), the composition of local groundwaters (large black circle), the range of calculated geothermal fluid compositions (shaded region) and the values of the vapor phase of those calculated aquifer fluids (open diamonds), and fractionation trends of vapor and remaining liquid at temperatures of 50, 100, 200, and 300 C, and liquid fractions of f = 0.9, 0.7, and 0.5. Liquid vapor fractionation from Horita & Wesolowski (1994). the low salinity of the condensate requires only minimal mass transfer from wallrock to fluid before the fluid reaches chemical equilibrium with respect to secondary minerals like epidote. In Fig. 7, isotopic fractionation trends between vapor and liquid of geothermal fluids with a starting composition equal to local groundwater (dd = 87&, d 18 O = 12.3&) are shown for temperatures between 100 and 300 C (thin dashed lines; Horita & Wesolowski 1994). Trends are shown for scenarios where the remaining liquid fraction is 50, 70, and 90% of the total starting fluid (f = 0.5, 0.7, 0.9). Epidotes from well K-20 (blue diamonds, Fig. 7) are consistent with precipitation from a fluid isotopically equivalent to the rising vapor phase of aquifer fluids, with the exception of the sample from 906 m depth (d 18 O = 9.7&), which appears to be in equilibrium with the liquid fraction of the aquifer fluid. Epidote from wells K-26 and K-34 (blue circles and squares, respectively) formed from fluids that slightly deviate from the fractionation trends of local meteoric water, possibly a result of the addition of volatiles exsolved from the magma heat source. Empirical and theoretical estimates of CO 2 :H 2 O ratios in basaltic magmas of ~1/15 (Sigvaldsson & Elısson 1968; Taylor 1986) indicate that 15% of Krafla geothermal waters are of magmatic origin, given CO 2 concentrations of wellhead fluids during the Krafla Fires. The result is local increases in dd FLUID of up to ~7& and d 18 O FLUID up to ~3.3& (Sheppard 1977). Geothermal fluids may further be affected by isotopic exchange with magmatic CO 2, but at observed geothermal temperatures, d 18 O fractionation approaches the existing isotopic variance in between local meteoric fluids ( 12&) and magmatic CO 2 (~ +6&), such that exchange between them would result in <1& variation (Bottinga 1968; Truesdell 1974). If geothermal fluids forming K-34 epidotes consisted of ~3 5% magmatic water, the data would approximate the corresponding f = 0.9 vapor or liquid fractionation curves in Fig. 7, at down-hole temperatures consistent with those approximated for this well (Table 5). Mixing of groundwater and magmatic gas followed by separation of the vapor phase also satisfies epidote isotope compositions from well K-26. This well is almost entirely composed of unaltered mafic intrusions, suggesting that it may be located close to a magmatic source, and thus enhanced boiling and enrichment in magmatic gases would be expected. While combined O and H isotope analyses of epidote are preferable for elucidating source(s) and isotopic evolution of fluids in individual wells of a geothermal system, inferences on hydrologic patterns such as flow direction and mixing patterns at Krafla are possible using only the dd of fluids in equilibrium with hydrothermal epidote. Figure 8 shows a contoured distribution of dd values for fluids in equilibrium with epidote along a northwest southeast trending cross section of the geothermal system (B-B in Fig. 1B). Black circles represent epidote sampling depths; the gray region approximates the lithologic boundary between the lower intrusive-dominated region and the upper region of hyaloclastites and lavas. Low dd values are generally concentrated along the lithologic boundary, suggesting that as deep geothermal fluids boil, the fractionated vapor phase rises upwards along vertical fractures until it reaches a highly permeable horizon at the lithologic boundary, along which fluids move laterally. The incursion of dd values < 94& in well K-17 may represent vapor rising along a southern upflow zone as identified by Armannsson et al. (1987) and Darling & Armannsson (1989) (Fig. 2B). High dd values concentrate along the Hveragil high-permeability zone suggesting that (i) geothermal fluids in this region have a greater component of magmatically sourced fluids, (ii) elevated temperatures at the base of Hveragil result in dd values of the vapor phase that are higher than the liquid phase (Horita & Wesolowski 1994), and (iii) high vertical permeability of this upflow zone possibly allows efficient transport of both liquid and vapor phases of the fluid (e.g., Coumou et al. 2008). The southeastward bend of high dd contours (upper part of K- 20) suggests that fluids migrating up Hveragil subsequently flow southward along the permeable lithologic boundary, feeding wells in the Sudurhlıdar region. Based on the chemistry of geothermal fluids and the spatially distinct stable isotope characteristics of geothermal epidote, in Fig. 9 we present a hydrogeologic model of the Krafla geothermal system that complements and expands

17 Krafla Hydrogeology 191 Fig. 8. Hydrogen isotope contours of water in equilibrium with hydrothermal epidote. Locations shown by black circles, along a NNW to SSE transect of the Krafla system (B B in Fig. 1B). Well K-34 is northeast of this transect, but projected onto the cross section to approximate fluid composition within the high-permeability zone, Hveragil. Approximate location of the lithologic boundary between porous hyaloclastites and basalts above and intrusives below is shown in gray. steam overlying the magma heat source, and injection of high-level dikes and sills in the caldera fill. Vapor rises rapidly along discrete steeply dipping faults and fractures in the Leirbotnar portion of the system, then mixes with cooler groundwaters of the upper aquifer. At Hveragil, increased permeability drives convective circulation along the length of the system, increasing temperatures of the primary aquifer to boiling, and allowing lower aquifer fluids to rise and flow along the upper aquifer located near the lithology transition. Fluids flowing upwards through Hveragil are vapor-rich, due to depressurization in the high-permeability zone, and CO 2 -rich due to increased input of magmatic volatiles. Southward migration along the hydrologic gradient of primary aquifer fluids mixed with vapor- and magmatic-rich lower aquifer fluids results in progressively lower dd FLUID values observed in Sudurhlıdar, Hvitholar, and Namafjall (Fig. 3). Epidote forming between the two aquifers occurs along the fractures where deeply derived vapors rise, cool and condense, and react with the surrounding host rock until reaching saturation with respect to epidote. upon previous interpretations (B odvarsson et al. 1984a; Giroud 2008; Arnorsson 2012). The primary source of geothermal fluids is local groundwater, sourced from highlands in the north. Groundwater flows along ~N S trending fissure swarms west of the geothermal system (Fig. 1), following a southeastern hydrologic gradient through the geothermal system and flowing in discrete permeable zones separated by low-permeability lithologies. Groundwater flow occurs primarily along the lithologic boundary between the upper sequence of hyaloclastites and lavas and the lower intrusive-dominated region. Less extensive horizontal flow occurs below ~1900 m, where groundwater is heated to boiling by a conductive layer of superheated CONCLUSIONS Phase segregation in wells of the Krafla geothermal system exhibiting excess enthalpy distorts the isotope composition of fluids discharged at the wellhead, such that they do not represent aquifer fluids. A closer approximation of aquifer fluid isotope compositions can be made by modeling the phase segregation process and estimating the original vapor fraction of the aquifer. Additionally, isotopic heterogeneity of fluids due to phase segregation is traceable throughout the geothermal system using the alteration mineral epidote. The spatial complexity observed in the temperature profiles and fluid chemistry of the Krafla geothermal system is Fig. 9. Updated model of hydrogeology in the Krafla geothermal system. All fluids have a meteoric source from the northeast. Deep fluids absorb mass and heat from the shallow magma chamber, and at fracture zones where pressure decreases, the geothermal fluid boils. Vapor rises into the shallow region of the system, mixes with colder groundwater and condenses. Both deep and shallow fluids continue into the Sudurhlıdar region, which has less distinctive, but still recognizable fluid stratification. B B refers to transect in Fig. 1B.