Scholarly article on topic 'The influence of oxygen isotope exchange between CO 2  and H 2 O in natural CO 2 -rich spring waters: Implications for geothermometry'

The influence of oxygen isotope exchange between CO 2 and H 2 O in natural CO 2 -rich spring waters: Implications for geothermometry Academic research paper on "Earth and related environmental sciences"

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Applied Geochemistry
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{"Oxygen isotopes" / "CO2 springs" / "Mineral springs" / Geothermometry / "Geochemical modelling" / "Water-rock reactions" / "Low temperature aqueous systems"}

Abstract of research paper on Earth and related environmental sciences, author of scientific article — Rūta Karolytė, Sascha Serno, Gareth Johnson, Stuart M.V. Gilfillan

Abstract Oxygen isotope ratio (δ18O) value deviations from the Meteoric Water Line with no significant change in the hydrogen isotope (δ2H) composition have been reported in naturally occurring CO2-rich waters from around the world. Here we review the effects of oxygen isotope exchange with CO2, high temperature equilibration with bedrock minerals and mineral dissolution and precipitation reactions on the CO2-rich water isotopic composition. We present two case studies from Daylesford (Australia) and Pah Tempe (Utah, USA) mineral springs, where we use a numerical geochemical modelling approach to resolve the influence of low temperature water-rock interactions and CO2 equilibration to the oxygen isotope ranges observed in the mineral waters. In both cases, we find that mineral dissolution – precipitation reactions are unlikely to have a significant effect on the groundwater isotopic compositions, and that the observed δ18O values in natural CO2 springs can be simply explained by equilibrium fractionation between water and free phase CO2. Traditionally, the interaction of CO2 and water in a natural CO2-rich groundwater setting has only been associated with water 18O depletion and this is the first study to consider 18O enrichment. We establish that in a natural setting, CO2 and water equilibration can result in water 18O depletion or enrichment, and that the change in the oxygen isotope composition ultimately depends on the initial CO2 and water δ18O values. Our new conceptual model therefore provides a mechanism to explain water 18O enrichment at ambient temperatures. This finding is critical for the use of δ18O in groundwater geothermometry and for the interpretation of natural water circulation depths: we argue that in some cases, natural waters previously interpreted as geothermal based on their oxygen isotope composition may actually have acquired their isotopic signature through interaction with CO2 at ambient temperatures.

Academic research paper on topic "The influence of oxygen isotope exchange between CO 2 and H 2 O in natural CO 2 -rich spring waters: Implications for geothermometry"

Accepted Manuscript

The influence of oxygen isotope exchange between CO2 and H2O in natural CO2-rich spring waters: Implications for geothermometry

Rûta Karolyté, Sascha Serno, Gareth Johnson, Stuart M.V. Gilfillan

PII: S0883-2927(17)30021-5

DOI: 10.1016/j.apgeochem.2017.06.012

Reference: AG 3914

To appear in: Applied Geochemistry

Received Date: 10 January 2017 Revised Date: 17 May 2017 Accepted Date: 22 June 2017

Please cite this article as: Karolyté, Rû., Serno, S., Johnson, G., Gilfillan, S.M.V., The influence of oxygen isotope exchange between CO2 and H2O in natural CO2-rich spring waters: Implications for geothermometry, Applied Geochemistry (2017), doi: 10.1016/j.apgeochem.2017.06.012.

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5180 %0 VSMOW

1 The influence of oxygen isotope exchange between CO2 and H2O in

2 natural CO2-rich spring waters: implications for geothermometry

4 Ruta Karolytè*1, Sascha Semo12, Gareth Johnson1, Stuart M.V. Gilfillan1

5 1School of GeoSciences, University of Edinburgh, Grant Institute, James Hutton Road,

6 Edinburgh, EH9 3FE, U.K.

7 Present address: School of Mechanical and Aerospace Engineering, Queen's

8 University, Ashby Building, 125 Stranmillis Road, Belfast, BT9 5AH, U.K.

9 *Author to whom correspondence should be addressed: ruta.karolyte@ed.ac.uk

10 Abstract

11 Oxygen isotope ratio (S18O) value deviations from the Meteoric Water Line with no significant change in the

12 hydrogen isotope (S2H) composition have been reported in naturally occurring CO2-rich waters from around

13 the world. Here we review the effects of oxygen isotope exchange with CO2, high temperature equilibration

14 with bedrock minerals and mineral dissolution and precipitation reactions on the CO2-rich water isotopic

15 composition. We present two case studies from Daylesford (Australia) and Pah Tempe (Utah, USA) mineral

16 springs, where we use a numerical geochemical modelling approach to resolve the influence of low

17 temperature water-rock interactions and CO2 equilibration on the observed oxygen isotope ranges observed in

18 the mineral waters. In both cases, we find that mineral dissolution - precipitation reactions are unlikely to

19 have a significant effect on the groundwater isotopic compositions, and that the observed S18O values in

20 natural CO2 springs can be simply explained by equilibrium fractionation between water and free phase CO2.

21 Traditionally, the interaction of CO2 and water in a natural CO2-rich groundwater setting has only been

22 associated with water 18O depletion and this is the first study to consider 18O enrichment. We establish that in

23 a natural setting, CO2 and water equilibration can result in water 18O depletion or enrichment, and that the

24 change in the oxygen isotope composition ultimately depends on the initial CO2 and water S18O values. Our

25 new conceptual model therefore provides a mechanism to explain water 18O enrichment at ambient

26 temperatures. This finding is critical for the use of S18O in groundwater geothermometry and for the

27 interpretation of natural water circulation depths: we argue that in some cases, natural waters previously

28 interpreted as geothermal based on their oxygen isotope composition may actually have acquired their isotopic

29 signature through interaction with CO2 at ambient temperatures.

30 Keywords: Oxygen isotopes; CO2-rich waters; mineral springs; geothermometry; geochemical modelling;

31 water-rock reactions; low temperature aqueous systems.

1. Introduction

The stable isotope ratios of water are a useful indicator of a wide range of geological conditions associated with groundwater reservoirs and mineral springs. Applications include interpretation of water origin (Harris et al., 1997; Ziegler, 2006; Demlie and Titus, 2015), residence times (Vuataz and Goff, 1986; Hearn et al., 1989;), migration pathways and mixing trends (Hearn et al., 1989; Siegel et al., 2004; Wilkinson et al., 2009; Delalande et al., 2011), temperature and circulation depths (Ziegler, 2006; Nelson et al., 2009), fault and fracture permeability to fluids (Cerón et al., 1998; Losh et al., 1999; Lions et al., 2014), local rainfall variations (Burns and Matter, 1995) and paleoclimates (Hays and Grossman, 1991). These geological interpretations rely on the identification of fundamental natural processes controlling kinetic and equilibrium fractionation of water stable oxygen and hydrogen isotopes (Fig. 1). Here we investigate the effect of water interaction with CO2 and its impact on the isotopic composition of natural waters. In particular, we investigate changes in the oxygen isotope ratio (S18O) values of water independent of changes in hydrogen isotope ratios (S2H).

The main processes that result in water 18O depletion or enrichment without affecting hydrogen isotope ratios are oxygen isotope exchange with CO2 and isotope exchange with bedrock minerals, either through low temperature mineral dissolution and precipitation, or diffusion at high temperatures (Fig. 1). Water 18O enrichment relative to the Meteoric Water Line (MWL) with no change in the hydrogen isotope ratio has been traditionally associated with oxygen isotope exchange with bedrock minerals or water-steam separation in geothermal conditions (e.g. Clayton & Steiner 1975; Matsuhisa et al. 1979; Giggenbach 1992). Due to the lack of other reported water 18O enriching mechanisms it has become common practice to interpret 18O enrichment in water as evidence for geothermal conditions (e.g. Cerón et al. 1998; Nelson et al., 2009). In contrast, oxygen isotope exchange with CO2 has been associated with water 18O depletion (Clark, I. D., Fritz, 1997; D'Amore and Panichi, 1987). However, recent applications of oxygen isotopes to monitor injected CO2 in Carbon Capture and Storage (CCS) reservoirs have demonstrated that the water isotopic composition after CO2 injection is primarily dependent on the oxygen isotope ratios of pre-injection water and CO2 and of the degree of CO2 saturation in the water and gas phases in the reservoir pore space (Kharaka et al., 2006; Johnson et al., 2011; Johnson & Mayer 2011; Serno et al., 2016). Applying this new knowledge to naturally occurring CO2-rich mineral springs, we re-evaluate previous interpretations of S18O and S2H geochemistry and propose a new conceptual model to explain the global occurrence of 18O shifts in CO2-rich waters.

In two case studies from Daylesford (Australia) and Pah Tempe (Utah, USA), we investigate the effects of mineral reactions and oxygen isotope exchange with CO2 to the observed water oxygen isotope composition. We propose a method to assess the relative contributions of these two mechanisms, which can be applied to waters with elevated CO2 concentrations in both natural and engineered settings.

2. Oxygen isotope geochemistry in spring waters

66 2.1 Global and Local Meteoric Water Lines

67 The ratios of stable oxygen (S18O = 18O/16O) and hydrogen (S2H = 2H/1H) isotopes in water are reported as

68 delta (S) notation relative to VSMOW (Vienna Standard Mean Ocean Water), following Equation 1, where R

69 represents the 18O/16O ratio for the sample and VSMOW, respectively.

'sample

I x 1000

71 There is a strong linear relationship between ô18O and S2H values in global precipitation reflected by the

72 Global Meteoric Water Line (GMWL), first defined by Craig (1961) and refined by Rozanski et al. (1993):

73 52H = 8.2 x 518O + 11.27

V-SMOW

74 The slope of the line is produced by equilibrium Rayleigh fractionation as precipitation is successively

75 removed from the vapour phase when it condenses, leaving the residual water vapour progressively depleted

76 in 18O and 2H. The intercept of the line is controlled by kinetic fractionation during evaporation of seawater.

77 Variations in humidity and temperature affect the slope and intercept of the line, and produce different Local

78 Meteoric Water Lines (LMWL) for specific locations (Dansgaard, 1964; Clark and Fritz, 1997).

79 2.2 Meteoric water stable isotope change due to natural processes

80 Various natural processes may deviate the stable isotope ratios in reservoir waters from equilibrium values on

81 the MWLs (Fig. 1). During evaporation, lighter isotopes enter the vapour phase, whereas in condensation

82 heavier isotopes are preferentially incorporated into the condensate. Consequently, water vapour is depleted in

83 18O and 2H whereas the remaining water is enriched. Similarly, temperature-dependent kinetic fractionation

84 occurs during steam loss above the boiling temperature, which produces 18O enrichment. (Clark and Fritz,

85 1997). Fractionation between degassing H2, H2S, CH4, and water in active magmatic systems leads to

86 enriched water S H values without an effect on S18O (Richet et al., 1977).

87 CO2-rich waters are often characterised by a horizontal deviation from the MWL (e.g. D'Amore & Panichi

88 1987; Pauwels et al. 1997; Cartwright et al. 2000). CO2 presence in the system can lead to water S18O changes

89 by either:

90 i) Promoting low temperature primary mineral dissolution and secondary mineral precipitation

91 reactions, preferentially consuming 18O.

ii) Equilibrium oxygen isotope exchange between CO2 and water.

93 Both of these processes affect only the water S18O values while S2H values remain unchanged, unless water-

94 rock reactions involve extensive precipitation of H-rich clays (D'Amore and Panichi, 1987). Additionally,

95 diffusive equilibrium oxygen isotope exchange with bedrock minerals at high temperatures produces 18O-

96 enriched waters. Here, we review and assess the relative contributions of these three processes (Fig. 1) to the

97 observed water S18O values of a global dataset of natural mineral waters.

5180 %0 VSMOW

99 Figure 1. Natural processes affecting water 618O and S2H values (adapted from D'Amore & Panichi, 1987).

100 2.2.1 Low temperature dissolution - precipitation reactions

101 The temperature-dependent oxygen isotope equilibrium fractionation factor between water and a precipitated

102 mineral leads to a preferentially 18O-enriched mineral phase and lower S18O values of the water. The extent of

103 this enrichment depends on the strength of mineral crystal lattice bonds (Zheng 2011). Under equilibrium

104 conditions, secondary minerals such as clays and carbonates preferentially incorporate more 18O during

105 precipitation and hence become enriched relative to the water (Compton et al., 1999; Kloppmann et al., 2002).

106 Clays also incorporate water molecules in the isotopically depleted intra layer, which leads to a positive S2H

107 value shift in the remaining pore water (Sheppard & Gilg, 1996). This process can significantly alter the water

108 S18O if the fraction of oxygen involved in the reactions is sufficiently high (D'Amore and Panichi, 1987).

109 Mineral dissolution and precipitation are particularly important in CO2-rich waters, which are often associated

110 with primary silicate hydrolysis and enhanced clay production rates (e.g., Watson et al. 2004; Kampman et al.

111 2014).

112 2.2.2. Diffusion related equilibrium oxygen isotope exchange with minerals

113 Meteoric water circulating at depth is depleted in 18O compared to rock-forming minerals. Oxygen isotope

114 exchange between the fluid and solid phases via diffusion gradually moves the two phases towards

115 equilibrium at a rate controlled by the water temperature. This is important in geothermal systems where the

116 disequilibrium is higher because the fractionation factor between water and minerals is low. In many cases,

117 deep groundwaters are enriched in 18O respective to the GMWL, with little change in S2H values as a

118 consequence of equilibrium oxygen isotope exchange with bedrock minerals that are commonly low in

119 hydrogen (Clark and Fritz, 1997). However, common rock-forming minerals such as feldspar, mica and quartz

120 require heating to temperatures above 250 °C to achieve this oxygen isotope exchange (D'Amore and Panichi,

121 1987). Hence the effect is only observed in geothermal conditions (Friedman and O'Niel, 1977).

122 2.2.3. Equilibrium oxygen isotope exchange with CO2

123 Equilibrium oxygen isotope fractionation occurs during redistribution of isotopes between two or more

124 compounds with forward and backward reactions proceeding at equal rates. At isotopic equilibrium heavier

125 isotopes preferentially concentrate in the phase with stronger bond constants (Young et al., 2002). In the case

126 of CO2 and water, CO2 has stronger bonds than water and thus equilibrium exchange results in the CO2 phase

127 being enriched in 18O. The fractionation factor (a) associated with equilibrium exchange reactions between

128 CO2 and water is expressed as:

130 where 518O = 18O/16O at equilibrium relative to VSMOW (Vienna Standard Mean Ocean Water). The

131 fractionation factor tfco2-H2o can be approximated as isotopic enrichment factor £co2-h2o expressed as a

132 difference between two reactants (in %o):

134 The fractionation factor is inversely correlated with temperature and is therefore particularly important for low

135 temperature waters (Bottinga, 1968). In most natural systems the oxygen isotope equilibrium between CO2

136 and water is predominantly influenced by the initial water S18O value as in the majority of low pressure

137 systems water represents the greater source of oxygen. However, in cases where CO2 represents a major

138 source of oxygen, the isotopic composition of water may be influenced by CO2 (e.g., Kharaka et al., 2006;

139 Johnson & Mayer 2011; Johnson et al., 2011; Serno et al., 2016).

129 aco2-H2o —

S1BOco2+ looo S18oH2o+ looo

133 103ln a ~ Sco2-H2o — Sl8Oco2 - 5l8Oi

2.3. Overview of global CO2-rich water isotopic compositions

Observed changes in the oxygen isotope geochemistry of global CO2-rich waters have been associated with isotopic equilibrium exchange between natural free phase CO2 and formation waters (Fig. 2). Examples from the literature include both thermal and cold springs with temperatures similar to those of ambient groundwater. Water 18O depletion and enrichment without a change in S2H has been observed in cold springs in the Valles Caldera-Southern Jemez Mountains in New Mexico, USA (Vuataz and Goff, 1986), along the Bongwana gas fault in South Africa (Harris et al., 1997) and in shallow boreholes in the Mont-Dore and Montmiral regions of the Massif Central, France (Casanova et al., 1999; Humez et al., 2014; Pauwels et al., 1997, 2007). The Alto Guadalentín groundwater aquifer in southeast Spain showed a similar shift in S18O which has been associated with the manifestation of CO2 from greater depth due to overexploitation of the aquifer, and oxygen isotope equilibrium exchange between the CO2 gas and reservoir water (Cerón et al., 1998; Cerón and Pulido-Bosch, 1999). Geochemical differences in various northern Portuguese CO2-rich mineral waters are reportedly caused by water-CO2 isotopic equilibrium exchange (Marques et al., 2000). A majority of the bubbling pools in the central Italian volcanic region, and some hydrothermal waters from Sicily, exhibit a S18O value deviation from the Local Meteoric Water Line (LMWL), which is likely the result of oxygen isotope exchange between the meteoric water and CO2 (Cinti et al., 2011).

A comparative study of thermal and cold spring waters in the Poroto-Rungwe region in Tanzania revealed that springs with observed sustained CO2 flux show water 18O depletion of up to -8 %o relative to the LMWL, while springs with relatively lower and episodic bubbling gas emanations are isotopically similar to surface meteoric water. Water 18O depletion was negatively correlated with temperature (Delalande et al., 2011), consistent with experimental observations indicating fractionation increase at low temperatures (Bottinga, 1968). Water 18O depletion caused by mantle-derived CO2 interaction with cold springs is also observed at the Sn^fellsnes Peninsula in Iceland. A number of hot springs from the same region are 18O enriched, although this is interpreted as a result of isotopic exchange with bedrock minerals at geothermal temperatures (Thomas et al., 2016). Finally, evidence from Hofsstadir (Iceland) shows that the effect of oxygen isotope exchange can be preserved in water after the gas has leaked out of the system. Stagnant geothermal water pools with low CO2 contents are reportedly depleted in 18O due to an episode of CO2 flux in the past (Kristmannsdóttir and Sveinbjoernsdóttir, 2012). Hence, this global compilation of observed CO2-rich water oxygen isotope composition clearly shows that 18O depletion or enrichment without a change in S2H is a common feature of low temperature CO2-rich springs and groundwaters and can be associated with both actively degassing and previously degassed systems. Strong horizontal trend in deviations from S18O values suggest this process is separate to the change in S2H values relative to current surface waters, which is attributed to recharge at different temperatures.

In our study, we focus on two natural examples of CO2-rich springs showing opposing linear S18O deviations from the MWL. CO2-rich mineral springs in Daylesford, south east Australia, show water stable isotope

175 values ranging from -7.8 to -5.8%o for S18O and -44 to -31.8 %0 for S2H, and are depleted in O relative to the

176 LMWL by up to 1.43 %o (Fig. 2b). Cartwright et al. (2002) attributed this to interaction with CO2 and

177 degassing at the surface. In contrast, Pah Tempe mineral waters in Utah, USA, range between 25 - 27.1%o for

178 S18O and -108.9 to -105.9 %0 for S2H, and are enriched in 18O compared to the LMWL by up to 1.68%0 (Fig.

179 2c).

180 The mechanism of water 18O enrichment at Pah Tempe is currently uncertain. Nelson et al. (2009) suggested

181 equilibrium isotope exchange with bedrock minerals at temperatures above 150 °C on the basis of the lack of

182 other plausible mechanism for water 18O enrichment. However, evidence for >3-5km deep faults providing

183 circulation pathways is limited (Nelson et al., 2009) and the temperature of the water discharging at the

184 surface ranges from 39 to 41 °C (Dutson, 2005). Further, circulation temperatures calculated from other

185 conventional geothermometry techniques indicate lower temperatures: 70-75 °C using conductive and

186 adiabatic silica and 37-39 °C using chalcedony silica geothermometers (Dutson, 2005) and up to 80 °C using

187 quartz and Na-K-Ca-Mg geothermometer (Budding & Sommer, 1986).

ö O %0 VSMOW -25 -20

-20 -40 -60 -80 -100

+ Daylesford (Cartwright et al., 2002)

> Massif Central (Pauwels et at, 1997)

s Montmirai, France (Pauwels et al., 1997)

♦ Alto Guadalentin, SE Spain (Cerón et al., 1998)

* Saratoga Springs, NY, USA (Siegel et al., 2004) o North Portugal (Marques et al., 2000)

■ Central Italy (Cinti et al., 2011)

■ Pah Tempe, Utah, USA (Nelson et al., 2009)

Ö ■ New Mexico (Vuataz and Goff, 1986) CO

> Tanzania (Delalande etal,, 2011) as

-170 T 4 Snsefellsnes Peninsula, Iceland (Thomas et al.,

È 2016)

• Hofsstadlr, Iceland (Kristmannsdóttir and Sveinbjoernsdóttlr, 2012)

5lsO %o VSMOW

-9 -8 -7 -6 -5 -4 -3 -2

• Daylesford springs (Cartwright et al., 2002) ■ Surface water (Cartwright et al., 2002)

5 O %o VSMOW -17 -16 -15

-10 -20 -30

Pah Tempe springs (Nelson et al., 2009) Surface water (Mayo et al., 2003)

-102 -104 -106 -108 -110 -112

-114 O S

-116«

-118 J -120-1

190 Figure 2. a) Global compilation of CO2-rich waters showing 18O depletion or enrichment without a change in ô2H

191 values relative to the GMWL. b) Daylesford: compilation of water isotopic composition in mineral spring waters,

192 previously published by Cartwright et al. (2002). ô18O values of spring waters are lower by as much as 1.43%o

193 relative to the LMWL. C) Pah Tempe: ô18O values of spring waters are higher than respective values on the

194 LMWL (Kendall and Coplen, 2001) by up to 1.68%« (data from Nelson et al., 2009). Full dataset is available as

195 supplementary table A.

196 3. Geological background of case studies

197 3.1 Daylesford springs, Australia

198 Daylesford mineral springs are located in the Central Highlands of Victoria, south east Australia (Fig. 3a).

199 There are more than 100 low temperature CO2-rich springs in the area, which have been exploited historically

200 for drinking and recreational purposes. Springs flow in a fracture-dominated aquifer through an Ordovician

201 turbidite sequence altered to greenschist facies and discharge into topographic lows such as streambeds. The

202 depth of circulation is unknown but historical records report spring water in mines at a maximum depth of 1.6

203 km (Shugg, 2009). The aquifer is overlain by Newer Volcanic basalts, active from 4.5 Ma to 5000 a (Boyce,

204 2013). The spring waters contain excess dissolved carbon (Weaver et al., 2006) and actively degasses at

205 surface discharge points. CO2 is reportedly mantle-sourced, based on their close proximity to the eruptive

206 centres (Lawrence, 1969), 3He/4He gas data (Chivas et al., 1983), and gas carbon isotope values (513C;

207 Cartwright et al., 2002). Spring waters are Na-HCO3 type with a significant solute composition variation

208 between individual springs (Weaver et al. 2006).

209 3.2 Pah Tempe springs, Utah, USA

210 The Pah Tempe springs discharge in the damage zone of the currently active Hurricane fault at Timpoweap

211 Canyon, Utah, USA (Fig. 3b). The Hurricane fault has a total displacement of up to 3000 m and a 200 m wide

212 core (Biek, 2003). All springs discharge at the eastern damage zone into the Virgin River Canyon. The

213 stratigraphic sequence includes Cretaceous sandstones, siltstones and shales, and carbonaceous Jurassic -

214 Triassic sediments and evaporites (Biek, 2003, Nelson et al., 2009). Free-phase CO2 is actively degassing at

215 the surface, forming prominent bubble trains, along with mineralised water at an elevated temperature of

216 40 °C. The spring waters are high in Na+, Cl-, Ca2+ and SO2- as well as dissolved CO2 (Dutson, 2005). Helium

217 isotope and S13C analysis indicate carbonate thermal alteration as a primary CO2 source (Nelson et al., 2009).

113° 15'60" 113" 16" 18"

| Quaternary basalts _ Fault

| | Quaternary sediments " Spring

I | Triassic mudstone, siltstone, sandstone, gypsium | | Permian gypsum, mudstone, limestone _ - _ Permian dolostone, gypsum, limestone, sandstone

220 Figure 3. a) Location map of Daylesford springs in Victoria, SE Australia. Springs emanate near major fault lines

221 and Newer Volcanic eruption centres (adapted from Cartwright et al., 2002). b) Pah Tempe springs in Utah, USA.

222 Springs discharge into the stream bed of Virgin River in a close proximity to the Hurricane fault (adapted from

223 Nelson et al., 2009).

4. Methods

4.1 CO2 sampling

CO2 samples from Daylesford and Pah Tempe springs were collected using a funnel placed over a bubbling vent and connected by plastic hosing to a copper tube. After purging air from the sampling line by allowing gas to flow for 5 to 10 minutes, the copper tubes were sealed by two metal clamps at either side creating a helium leak tight cold weld. At Daylesford water samples were collected via hand pumping the mineral water wells and from stream waters at gas discharge locations where CO2 samples were collected. Waters were filtered through 0.45 ^m pore-size filters and filled into Nalgene bottles (with no headspace) or vacutainers. Samples were stored in a cooler until analysis to avoid evaporation. The temperature of the water which the CO2 was bubbling through was measured in the field using a Hannah Instruments HI991300 Portable Waterproof pH/EC/TDS Meter. Temperature accuracy is ±0.5°C.

4.2 Stable isotope analysis

The S18O values of exsolved CO2 gas samples from Pah Tempe and Daylesford were measured at the Scottish Universities Environmental Research Centre (SUERC). Gas samples were released into an ultra-high vacuum extraction line. Two aliquots of the gas were trapped into glass ampules for stable isotope analysis, using liquid nitrogen and sealing the ampule with a blowtorch after pumping away the other volatile gases. Samples were analysed using a VG Isotech Sira 2b mass spectrometer with typical uncertainties of ±0.3%o. S18O and S2H measurements of two Daylesford water samples were obtained at the University of Wollongong, School of Earth and Environmental Sciences Stable Isotope laboratory using a Micromass PRISM III mass spectrometer. Uncertainties for S18O and S2H measurements are ±0.1%o and ±1%o respectively.

4.3 Numerical simulation of water-rock reactions

Numerical simulations allow the quantification of the mass transfer between the solid and fluid phases during dissolution and precipitation reactions at elevated CO2 pressures. The amount of oxygen liberated from dissolving minerals and precipitated in secondary minerals can be compared to the total oxygen in the solution to assess the contribution of water-rock reactions to the overall water oxygen isotope signature.

Mineral dissolution and precipitation reactions are simulated using the geochemical modelling software PHREEQC (Parkhurst and Appelo, 1999) with the WATEQ4F database (Ball and Nordstrom, 1991). The numerical simulations solve a set of nonlinear mass balance equations using thermodynamic constants defined in the database.

Simulations require a set of reactive primary and precipitating secondary minerals, which are chosen according to the saturation indices of the dissolved species in groundwater calculated in PHREEQC and the

geology of the local areas. The main goal of the simulations was to reproduce the observed geochemistry of the spring waters, published in Weaver et al. (2006) for Daylesford springs and Dutson (2005) for Pah Tempe springs, and to quantify the oxygen isotope transfer between the solid and fluid phases. There are multiple reaction pathways to achieve the observed water chemistry but the model seeks to represent the simplest solution using the most likely reactive phases based on the geological setting.

4.3.1 Model assumptions

4.3.1.1 Initial groundwater and CO2 dissolution

Equilibrium dissolution and precipitation reactions are modelled in a closed system considering reactive species in 1 litre of water. Average modern local groundwater is used as a starting solution (Table 1). CO2 partial pressure (pCO2) is calculated from measured pH and alkalinity (as HCO3-) values, so mineral reactions are modelled at ambient pressures and average recorded spring discharge temperatures (25 and 40 °C for Daylesford and Pah Tempe, respectively). This method provides the best fit to the measured DIC contents. Alternatively, reactions could be modelled at depth with higher pCO2 and subsequent dilution at the surface. This would allow dissolution of more stable minerals in the beginning (May, 2005) but was found not to have a significant effect on the overall mass balance.

pH Total Alk. meq/L Cl- mg/ L SO42-mg/L Ca2+ mg/L K+ mg/L Mg2+ mg/L Na+ mg/L

Daylesford 5.85 0.12 5.36 1.76 1.14 0.48 0.42 3.11

Pah Tempe 7.01 0.88 23.3 2.3 81.7 3 20 25.6

Table 1. Baseline water compositions. Daylesford: weighted average Melbourne precipitation water between 2007 and 2011 (Crosbie et al., 2012). Pah Tempe: average Virgin River valley groundwater 2015 (Burden, 2015).

4.3.1.2 Mineral reactions in Daylesford, Australia

Major cation concentrations and bicarbonate contents display a positive correlation, indicating that mineral dissolution is proportional to acid neutralisation This is a common feature in Na-Ca-HCO3 type waters produced by acid groundwater dissolution of silicates (May, 2005). These trends infer that the bulk reaction is limited by the kinetics of primary mineral dissolution and that the system is not in equilibrium. SiO2 does not show correlation with HCO3-, suggesting secondary silica precipitation.

The model assumes reactions with common minerals in the Ordovician Castlemaine turbidite sequence and Newer Volcanics intrusions. Albite and chlorite were chosen as source rocks for Na, Mg and Fe in the water based on XRD analysis of average Ordovician Castlemaine turbidites collected in the Ballarat area (Bierlein et al., 1999). The study compares unaltered Ordovician rock with that altered by hydrothermal fluids. Two other major phases, muscovite and quartz, are stable at hydrothermal conditions, and therefore are assumed to be unreactive in CO2-water system in the model. Ca and small amounts of K are sourced from feldspars in mafic Ca-rich and trachytic lavas from the Newer Volcanic sequence (Price et al., 2003), and are modelled as anorthite and adularia. Dissolution of minor amounts of sulfates (melanterite, barite) contributes SO4, Fe and

286 Ba to the system, consistent with redox values measured by Weaver et al. (2006). Other trace elements (Sr,

287 Mn) are sourced from carbonate dissolution (rhodochrosite, siderite). Secondary minerals are allowed to

288 precipitate to equilibrium were kaolinite, amorphous silica and Mg-carbonates.

289 4.3.1.3. Mineral reactions in Pah Tempe, Utah, USA

290 The main model assumptions are based on spring water geochemistry interpretation by Dutson (2005). There

291 is little variation in solute contents between samples collected at different discharge sites (Dutson, 2005),

292 suggesting springs emerge from a single aquifer. The average molar ratios for Na/Cl and Ca/SO4 are 1.08 and

293 0.98, indicating halite and gypsum dissolution, which could be sourced from the Triassic sequence (Biek,

294 2003; Dutson, 2005). As gypsum accounts for most of the Ca contents, carbonate dissolution is unlikely.

295 Minor amounts of Mg and K are introduced by silicate dissolution, modelled as phlogopite. Precipitation of

296 amorphous silica provides a sink for Si.

5. Results

5.1 Stable isotope composition

Oxygen isotope ratios of CO2 (S18Oc02) degassing at the surface of spring discharge points in Daylesford are 36.3%o and 34.1%o. The water sampled near gas flux points has S18O values of -6.3%o and -5.55%o and S2H values of -34.6%o and -33.1%o. S18Oc02 values of two Pah Tempe springs are 25%o and 26.8%o (Table 2).

Location Spring ID S180H2o %« VSMOW S2H VSMOW S#$0Co2%o VSMOW T °C

Daylesford Locarno -6.30 -34.6 36.3 16.7

Daylesford Taradale -5.55 -33.1 34.1 20.9

Pah Tempe Virgin No4.7110106 25

Pah Tempe PAH TEMPE 1.B 26.8

Table 2. S#$0 Co2, S180H20, S 2H and temperature values for mineral water gas discharges collected in Daylesford and Pah Tempe.

5.2 Low temperature mineral precipitation - dissolution reactions

Numerical simulations of equilibrium dissolution and precipitation of primary and secondary phases at fixed CO2 partial pressures (0.8 atm in Daylesford and 0.79 atm in Pah Tempe) produce geochemical compositions that closely match the reported measurements of Na-HCO3 and Na-Cl-HCO3 waters from Daylesford and Pah Tempe (Fig. 4). In Daylesford, the modelled Na, Ca, K, SO4 and trace element (Fe, Sr, Mn and Ba) contents are fixed by the defined amount of primary dissolving minerals, while Mg and Si concentrations are controlled by equilibrium secondary mineral precipitation. For Pah Tempe springs, Na, Cl, Mg and K contents are controlled by fixed amounts of mineral dissolution, Ca and SO4 are limited by the gypsum solubility, and Si is controlled by the kaolinite and secondary silica precipitation. Both models overestimate Si contents. The exact precipitating silica polymorphs are unknown and the thermodynamic constant for amorphous silica used in the simulation may not be precise. The relative amounts of dissolving minerals are not proportional to the bulk rock composition (as reported by Bierlein et al., 1999) suggesting that carbonic acid alteration is limited by mineral dissolution kinetics. The modelling approach for which fixed amounts of minerals are dissolved to match the observed element concentrations effectively eliminates the uncertainties associated with predicting individual mineral dissolution rates.

The total amount of phases that react are summarised in Table 3. In both cases, mineral reactions liberate more oxygen to the system than remove via precipitation. However, the total amount of oxygen involved in both types of reactions represents only 0.11 % and 0.087 % of total oxygen atoms in the water in Daylesford and Pah Tempe, respectively.

TotalC Na Ca Mg K Si SCU Fe Sr Mri Ba

324 Figure 4. Geochemical modelling results compared to average concentrations from the literature. a)

325 Molality in solution at 25 °C in Daylesford b) Pah Tempe springs (40 °C). Black dots represent

326 measured values reported by Weaver et al. (2006) for Daylesford and Dutson (2005) for Pah Tempe

327 springs. The dashed black line is an average of the measured values. The blue solid line shows modelled

328 element concentrations.

Dissolved Precipitated O dissolved O precipitated

Phase mol/L mol/L mol/L mol/L

Daylesford springs

Adularia KAlSi3O8 3.00E-04 1.85E-04

Albite NaAlSi3O8. 1.20E-02 7.38E-03

Anorthite CaAl2Si2O8 4.00E-03 2.46E-03

Barite BaSO4 2.00E-06 1.33E-06

Celestite SrSO4 1.00E-05 6.67E-06

Chlorite (Mg,Fe)3(Si,Al)4O1o 9.00E-04 3.75E-04

CO2(g) CO2 4.67E-02 3.11E-02

Melanterite FeSO4-7H2O 6.00E-05 2.44E-05

Rhodhochrosite MnCO3 6.00E-06 3.60E-06

Dolomite CaMg(CO3)2 1.31E-03 7.85E-04

Kaolinite Al2Si2O5(OH)4 1.11E-02 5.85E-03

Amorphous silica SiO2 2.36E-02 1.57E-02

Total 0.06 0.04 0.04 0.02

% of total O in 1 L of water 0.07 0.04

Pah Tempe springs

CO2(g) CO2 5.54E-02 > 3.69E-02

Gypsum CaSO4-2H2O 1.91E-02 9.53E-03

Halite NaCl ^s1.00E-01 0

Phlogopite KMg3(AlSi3Oio)(F,OH)2 2.00E-03 1.00E-03

Amorphous silica SiO2 2.39E-03 1.59E-03

Total 0.18 0.002 0.047 0.002

% of total O in 1 L of water 0.085 0.003

329 Table 3. Summary of dissolving and precipitating minerals in Daylesford and Pah Tempe springs, and

330 associated oxygen concentrations contributing to the water. Oxygen sourced from mineral reactions is

331 compared to the total oxygen atoms in 1 litre of water (55.6 mol/L).

332 6. Discussion

333 6.1 Water - rock reaction influence to water oxygen isotope composition

334 The effect of mineral dissolution and precipitation on the water oxygen isotope composition depends on the

335 S18O values of the water, dissolving and precipitating phases, and the relative ratios between the solid and

336 fluid phases. This relationship can be expressed by a simple mass balance model:

337 X( X 51S0( + X.X 51S0. = X(X 5180( + X. X 51S0. (5)

338 Where X is the relative fraction of oxygen in the phase, and 51S0l and 51S0f are the initial and final

339 oxygen isotope ratios in mineral (m) and water (w). The mass balances obtained from numerical simulations

340 (Table 3) indicate that mineral reactions account for dissolution of 0.7 g and 0.8 g, and precipitation of 0.4 g

341 and 0.04 g of oxygen per 1 litre of water in Daylesford and Pah Tempe, respectively. This represents ~ 0.1%

342 of the total oxygen in the water. The theoretical water S18O change in Daylesford and Pah Tempe given a S18O

343 range of viable rock forming mineral values (5 - 40%o) does not exceed 0.01%o, which is below the analytical

344 sensitivity and therefore has a negligible influence on the S18O.

345 It is also important to consider mineral reactions that achieve equilibrium quickly and may not be evident in

346 the water geochemistry, such as primary carbonate dissolution and secondary carbonate precipitation. A

347 theoretical scenario of carbonate dissolution with low S18O values formed at high temperatures and re-

348 precipitation at low temperatures would result in a water 18O depletion of 0.5 %o, requiring progressive

349 reworking of 200 g of carbonates per 1 litre of water. A recent study by Sterpenich et al. (2009) demonstrated

350 that less than 1% by mass of an oolitic limestone dissolved due to interaction with CO2-saturated water under

351 extreme experimental conditions (150 bar, 80 °C). These results clearly show that carbonate dissolution and

352 re-precipitation at the amounts required to significantly alter S18O values is unlikely.

353 The mineral water geochemistry in Daylesford and Pah Tempe is controlled by the primary mineral

354 dissolution and secondary mineral precipitation accelerated by elevated CO2 partial pressures. However, the

355 fraction of oxygen in these reactions is too small to influence the oxygen isotope ratio of the water body.

356 Consequently, water-rock reactions cannot account for water S18O deviations from the LMWL in both case

357 studies due to their low salinity. However, the method could be applicable to more saline formations where

358 mineral reaction may liberate enough oxygen to alter water S18O values, such as deep basement fluids and

359 hypersaline brines. Depletion of 18O in saline brines (up to 250g/L) in Fennoscandian and Canadian Shields

360 have been previously interpreted as a result of low temperature water-rock reactions (Blomqvist, 1990; Frape

361 and Fritz, 1982).

6.2 CO2-water oxygen isotope exchange influence on water isotopic composition

Our geochemical modelling results clearly show that low temperature mineral reactions can be excluded as a significant source of oxygen to the waters at Daylesford and Pah Tempe. Hence these two case studies, along with the global compilation of stable isotope values from CO2-rich springs, provide robust evidence that CO2-water oxygen isotope equilibrium exchange in the subsurface can result in 18O depletion and enrichment in CO2-rich spring waters compared to the MWL.

The amount of CO2 required to achieve the S18OH20 change observed in Daylesford (-1.7%o) and Pah Tempe (2 %o) can be estimated using the conceptual model developed by Johnson et al. (2011). The magnitude of the shift in S18OH20 relates to the fraction of CO2 in the system. The extent to which CO2 can change the oxygen isotope composition of reservoir water depends on the:

• Initial S18Oc02

• Initial water S18O value (S18OH20) calculated from the LMWL)

• Relative proportions of CO2 and water that equilibrate (X)02 ) as the fraction of oxygen sourced from CO2 in the system)

• Temperature-dependent oxygen isotope enrichment factor (£C02~H2o)

This relationship is expressed in equation (6) where 18O^20 is the final oxygen isotope composition of water (Johnson et al. 2011):

ô1QO/20 = (Ô1QOC02 - ÏC02-H20) X X002 + S18OlH20 X(l- X002) (6)

The CO2 source determines the initial 518OC02, which is an important control on the ultimate water S18O shift achieved at equilibration. Generally, S18O values in rocks decrease with increasing formation temperature (Fig. 5). Due to this difference of initial values, mantle CO2 has the potential to produce the most 18O-depleted fluids after equilibration, while CO2 generated through thermal carbonate alteration may enrich water in 18O.

In Daylesford we consider two examples of &18OC02 values associated with mantle degassing. The first one (8%o) represents a 'minimum value' scenario of CO2 degassed from a volcanic source whereby large amounts of mantle-sourced gas ascends through the crust without significant interaction with other fluids than the mineral springs. However, S18O is likely to change due interaction with subsurface fluids with increasing migration or residence time in a natural trap. The second scenario represents mantle CO2 after interaction with subsurface fluids and uses S18O value of 19.1%o measured in Caroline CO2 field in Mt.Gambier, South Australia (Chivas et al., 1987), which is associated with the same period of volcanic activity as the eruptive centres in Daylesford. A value of 518OC02 of 30%o represents an average value for thermogenic CO2 from a

393 global range (Bindeman, 2008). Equilibration with these potential values for mantle-derived CO2 at

394 Daylesford would require the fraction of oxygen sourced from CO2 in the system ('Co2) to be 5% and 10%

395 for average mantle and Caroline field CO2, respectively, to explain the maximum observed S18O deviations.

396 The maximum S18O shift observed in Pah Tempe can be explained by equilibration with the average

397 thermogenic CO2 when ')o2 is 30%. This amounts to 60 - 120 g and 366 g of CO2 per every litre of water in

398 Daylesford and Pah Tempe, respectively.

399 This simple model uses a closed system two-component mixing approach. In reality, both CO2 and water

400 move through the system at different rates. If CO2 moves through a relatively stagnant water body at a

401 continuous rate and degasses at the surface, the calculated 'co2 ranges represent the amount of CO2 that the

402 water has interacted with, rather than the amount of CO2 currently present in the system. Therefore, these

403 values can be taken as a maximum estimate.

404 The variation of oxygen isotope ratios in MORB-type igneous rocks are between 5.5 and 5.9%o, rhyolitic

405 magmas have values between 5.8 - 6.5%o and the overall range of various measured igneous lithologies is 4

406 to 12%o (Figure 5 inset) (Bindeman, 2008). CO2-mineral isotopic enrichment factors measured at melting

407 temperatures range between 2 - 6%o (Matthews et al., 1994; Stolper and Epstein, 1991), giving the overall

408 range of volcanic degassing related 518O4o2 values of 6 - 18%o with MORB-like signatures at the lower end

409 of the spectrum. In contrast, S18O values of sedimentary and carbonate rocks range between 8 to 32%o. CO2-

410 mineral isotopic enrichment factors span from 2 to 11%o depending on temperature and are at the higher end

411 of this spectrum in temperatures relevant to thermogenic gas generation. CO2-mineral isotopic enrichment

412 factors span from 2 to 11%o depending on temperature and are at the higher end of this spectrum in

413 temperatures relevant to thermogenic gas generation (Zheng, 1999). Considering the fact that meteoric water

414 S18O values span from -65 to 0%o and the wide range of natural 518OC02 values, it is clear that oxygen isotope

415 exchange with CO2 can both deplete or enrich water in 18O within the range of naturally occurring 518Olc02

416 and S18OH20 values, as demonstrated using the case studies from Daylesford and Pah Tempe.

5180(H20)%o VSMOW -15 -10

5180(C02) %o VSMOW

o 8 19.1 30

Mantle CO2

•30 *

Caroline CO2 gas field, S Australia (Chivas et al , 1987)

■50 *

'CO^oyo

Thermogenic CO2

Basalts Andesites Rhyolites Carbonates and diatoms Pelagic clays Sandstones and clastic Metamorphlc rocks Meteoric water

5 10 20

5lsO%o VSMOW

adapted from (Bindemai 2008)

• Daylesford springs, Australia (Cartwright et al., 2002)

• Pah Tempe springs, Utah, USA (Nelson et al., 2009)

• 5180(C02) inital values

Xco, Fraction of O sourced from CO2 In the system (%)

Figure 5. Water samples from Daylesford (blue) and Pah Tempe (green) showing 618O deviations from the

GMWL. The additional graph on the right shows potential initial o 0C02 values the water could have equilibrated with. The dashed lines with percentages show the X°C02 required to produce the observed shift for chosen examples. Average mineral S18O values provided in the inset at the bottom left (adapted from Bindeman, 2008).

423 6.3 Differences between observed and theoretical enrichment factors

We combine newly obtained gas S1oOC02 and water S1oOH20 measurements with those previously published in Dutson (2005) to calculate the expected oxygen isotope enrichment factor (£co2-h2o

) (Bottinga, 1968) for

individual springs (Table 4). For Pah Tempe springs calculations, we use an average temperature for springs where data is not available and S180C02 value of 26.3 ± 0.9 %o, which is an average of two measurements reported in this paper and one measurement of 27.1%o reported in Dutson (2005). Temperature-dependent oxygen isotope enrichment factors £C02-H20 calculated for the individual springs are larger than the observed difference between S 18O of water and CO2 (A) in all cases except for Taradale spring in Daylesford. The difference between the observed and theoretical enrichment factor (A -s) is 0.2%o and -1.9%o for two Daylesford springs and 2 ± 0.4 %o for Pah Tempe. Potential reasons for this apparent dis-equilibrium include the effects of salinity, partial equilibrium during CO2 ascent and kinetic fractionation on bubble surfaces.

These mechanisms have opposing isotopic effects: bubble formation on the surface leads to lower apparent CO2-water fractionation factor, while salinity and partial equilibration can either decrease or increase it.

Sample Temperature S2H ™ofh 0 smoc02 180f20 %o £co2-h2o A %« A - <

ID °C %» %o" %„ %« %

Pah Tempe springs, Utah, USA

S 7 40.6 -108.0 -13.0 26.3 ± 0.9 -14.2 37.2 39.0 1.8

S8 40.9 -108.3 -13.6 26.3 ± 0.9 -14.3 37.1 39.7 2.6

S 9 39.4 -107.7 -13.9 26.3 ± 0.9 -14.2 37.4 40.0 2.6

S 10 40 -108.7 -13.6 26.3 ± 0.9 -14.3 37.3 39.7 2.4

S 11 40 -108.5 -13.5 26.3 ± 0.9 -14.3 37.3 39.6 2.3

S 12 40.2 ± 0.6 -108.3 -13.2 26.3 ± 0.9 -14.3 37.2 39.2 2.0

S 13 40.2 ± 0.6 -108.4 -13.2 26.3 ± 0.9 -14.3 /<,37.2 39.2 2.0

S 14 40.2 ± 0.6 -105.9 -12.8 26.3 ± 0.9 -13.9 37.2 38.8 1.6

S 15 40.2 ± 0.6 -108.3 -12.9 26.3 ± 0.9 -14.3 37.2 38.9 1.7

S 16 40.2 ± 0.6 -108.3 -12.8 26.3 ± 0.9 -14.3 37.2 38.8 1.6

S 17 40.2 ± 0.6 -108.6 -13.1 26.3 ± 0.9 -14.3 37.2 39.1 1.9

S 18 40.2 ± 0.6 -107.4 -13.7 26.3 ± 0.9 -14.1 37.2 39.8 2.6

S 19 40.2 ± 0.6 -108.9 -12.7 26.3 ± 0.9 -14.4 37.2 38.7 1.5

S 20 40.2 ± 0.6 -108.6 -13.0 26.3 ± 0.9 -14.3 37.2 39.0 1.8

S 21 40.2 ± 0.6 -108.2 -13.1 26.3 ± 0.9 -14.3 37.2 39.1 1.9

S 22 40.2 ± 0.6 -108.5 -13.6 26.3 ± 0.9 -14.3 37.2 39.7 2.5

S 23 40.2 ± 0.6 -108.4 -12.8 26.3 ± 0.9 -14.3 37.2 38.8 1.6

Daylesford springs, SE Australia

Locarno 16.7 -34.58 -6.30 ^ 36.4 -5.4 41.9 42.1 0.2

spring-3

Taradale 20.9 -33.10 A-5.55 34.1 -5.2 41.0 39.1 -1.9

spring-1

Table 4. Summary of temperature and isotope measurement data in Pah Tempe and Daylesford springs used to calculate theoretical (£C02-f20) and observed enrichment factors (A k S18Oc02 - S180f20). The observed enrichment factor A k S180C02 - S180f20 is calculated as A = 1000lna; a = (S180C02+1000)/( S180f20 + 1000). All temperature and water isotope data from Pah Tempe springs from Dutson (2005). S180C02 value for Pah Tempe springs an average from 2 measurements reported in this paper (Table 2) and one measurement (27.1%) from Dutson (2005). LMWL equations used to calculate 180f20 % as follows: Daylesford 82H = 8.27 x 818O + 10 (Liu et al., 2010); Pah Tempe 82H = 6.7 x 818O - 12.6 (Kendall and Coplen, 2001).

6.3.1 Water salinity

High salinity waters display slightly altered £co2-h2o depending on temperature and total dissolved solids (TDS). Truesdell (1974) and Becker et al. (2015) reported £c02-h20 decreasing by 1%o in laboratory experiments using NaCl solutions of up to 250 g/L. Lecuyer et al. (2009) reported an increase in fractionation factor by up to 0.5 %o for 250 g/L KCl and sea salt solutions, which highlights the fact that different types of

ions may have opposing effects. Given the relatively low salinities at Pah Tempe (8 g/L) (Dutson, 2005) and Daylesford (5 g/L) (Cartwright et al., 2002), this effect should be negligible.

6.3.2 Partial equilibration during CO2 ascent to the surface

CO2-water isotopic equilibrium established locally or at a certain depth may be affected by the kinetics of two-phase fluid flow as CO2 ascends to the surface. The extent of mixing achieved by CO2 and water depends on the nature of interaction at depth. The observed water S18O shift requires large quantities of free phase CO2 interacting with the water. There is significant heterogeneity associated with the two-phase CO2 and water flow through the subsurface due to pressure and temperature effects on CO2 physical properties. CO2 and water may interact as transient, dispersed or separated two-phase flows at varying rates, depending on bedrock properties and interfacial tensions between the two phases (Plampin et al. 2014; Roberts et al. 2014). As pressure and temperature decrease at shallow levels of the subsurface, increased CO2 buoyancy provides a driving force to migrate at a faster rate. CO2 may achieve local equilibrium with surrounding water at depth where the flow rate is relatively low and where the isotopic signature is preserved after rapid ascent to the surface. Consequently, the remaining water will re-equilibrate and display a smaller S 18O shift than expected if in equilibrium with the sampled CO2. This may be the case in Pah Tempe springs where the CO2 flux is strong and sustained - measured S1oOC02 may represent equilibrium with water enriched in 18O by up to 2%o more than measured. Degassing in Daylesford is much more diffuse and episodic, allowing more time for equilibration with surface water.

6.3.3 Kinetic fractionation on bubbles

An alternative or additional mechanism, which may deviate S1oOC02 from equilibrium, is kinetic fractionation during diffusion of dissolved species towards gas bubbles at the surface. Mass control on diffusivity leads to preferential uptake of 12C and 16O during bubble formation (Affek & Zaarur 2014). If time is not sufficient for re-equilibration, the fractionation factor between CO2 and water will be lower than expected. In both localities gas samples were collected at the surface of bubbling streams and may therefore represent a lower-than-equilibrium S18OC02 value.

Kinetic fractionation on bubbles has been extensively studied in relation to the volatile content in degassing volcanic melts (e.g. Aubaud et al. 2004; Paonita & Martelli 2006). The effect has been observed as S13C deviations of dissolved inorganic carbon (DIC) from equilibrium by up to 4 %o in groundwater springs, seepage waters and headwater catchments (Doctor et al. 2008), and up to 2.5 %o deviation from equilibrium in DIC S13C samples collected at cold water springs in Green River, Utah (Assayag et al., 2009). The extent of kinetic fractionation increases with water pH and decreases with the depth of degassing. Liquids with lower volatile supersaturation are reported to show lower kinetic fractionation effects as equilibrium can be reestablished quicker (Affek & Zaarur 2014).

The limited availability of S18Oc02 measurements does not allow a quantitative comparison between the deviation from equilibrium in springs relative to measured pH and DIC contents. Generally, more actively degassing springs would be expected to deviate more from equilibrium, which agrees with field observations of higher gas discharge rates at Pah Tempe relative to Daylesford.

6.3.4 Summary

Water and CO2 sampled at the surface is unlikely to be in equilibrium due to localised partial equilibration and kinetic fractionation at the surface, while salinity is unlikely to have an effect. Values of S18Oc02 from high flux springs in Pah Tempe and Taradale spring in Daylesford suggest that the oxygen isotope signature from localised equilibration at depth is preserved. CO2 from diffuse Daylesford springs (e.g. Locarno) reequilibrates with dilute shallow water. The S18Oc02 values of gas collected at the water surface may be further obscured by 18O depletion during bubble formation, which affects high flux springs more than those with a diffuse low flux.

6.4 Implications for usage of 518O values in geothermometry

Water 18O enrichment relative to the MWL with no effect on S2H has been traditionally associated with geothermal systems. Waters enriched in 18O are produced by isotopic exchange between hydrothermal fluids and bedrock minerals, normally at temperatures above 250 °C (Clark, I. D., Fritz, 1997; D'Amore and Panichi, 1987). The fractionation factor between any mineral and fluid is governed by temperature, thus allowing the distribution of isotopes to be used as a geothermometer (e.g. Giggenbach, 1992). Another mechanism is water-steam separation above liquid-vapour isotopic exchange at 220 °C. At this temperature oxygen is kinetically fractionated between fluid and vapour but there is no fractionation in hydrogen isotopes (Clark, I. D., Fritz, 1997). Due to the lack of other reported water 18O enriching mechanisms, it has become common practice to interpret 18O enrichment in reservoir water as evidence for geothermal conditions (Ceron et al., 1998; Nelson et al., 2009).

Here, we present evidence that oxygen isotope exchange with CO2 can result in 18O-enriched waters, if the starting S18O value of CO2 is significantly higher than that of the water and if high gas/to water ratios are present. The oxygen isotope exchange between CO2 and the spring waters provides a more robust explanation for the 18O-enriched waters from Pah Tempe springs, which is in closer agreement with the geothermometric calculations and water discharge temperatures without invoking circulation depths of over 5 km and equilibration at temperatures >150 °C as quoted in Nelson et al (2009).

Our findings have two significant implications. Firstly, CO2-water equilibration alone, without the need to invoke any additional processes, may result in water O enrichment or depletion, which means that 18O-enriched waters should not be solely interpreted as geothermal as is current practice. Secondly, equilibrium

achieved between water and minerals at depth may be obscured by later interaction with CO2. Both of these scenarios have significant implications to the sulfate-water oxygen isotope geothermometry technique, which relies on temperature and pH-dependant oxygen isotope exchange between water and dissolved SO4, applicable to temperature ranges between 140 and 350 °C (McKenzie and Truesdell, 1977). The method requires estimation of the water S18O in equilibrium with sampled sulfate under an assumption that the original value has not been altered by secondary processes such as dilution with shallow water, boiling and steam loss, near-surface oxidation of H2S and biological activity or an application of an appropriate correction (Fowler et al., 2013). Equilibration with CO2, which can be achieved in a matter of hours and either deplete or enrich water in 18O, should also be considered when using this geothermometry technique in CO2-rich waters. This consideration may also be significant to palaeowater studies which relate S18O values of precipitating phases to either the palaeowater composition or precipitation temperatures (e.g. Astin and Scotchman, 1988; Morad and Eshete, 1990).'

A recent study by Ladd & Ryan (2016) demonstrated that shallow surface build up in CO2 partial pressure and subsequent bubble formation may be the main driving mechanism for geyser eruption in sub-boiling conditions, challenging the common notion that subsurface water boiling is required for this phenomena. Our study provides additional evidence that elevated CO2 concentrations at ambient temperatures may explain the features often attributed to geothermal systems.

7. Conclusions

Global natural CO2-rich mineral waters show S18O deviations from the MWL with no observed change in S2H. Oxygen isotope deviations without a change in hydrogen isotopes may be the result of oxygen isotope equilibrium exchange between CO2 and water, mineral dissolution and re-precipitation, or isotopic exchange with minerals. We have developed a simple geochemical modelling approach to study the influence of low temperature water-rock reactions on oxygen isotope changes in subsurface waters. The method requires knowledge of the water geochemistry (major ion concentrations, dissolved carbon content, pH, temperature) and a conceptual model of reactive and precipitating phases. Numerical modelling can be applied to assess the water-rock interaction influence on oxygen isotope ratios in other saline natural waters or CO2 storage sites where oxygen isotopes are used as natural tracer of the injected CO2 plume.

In two case studies from Daylesford (Australia) and Pah Tempe (Utah, USA), we apply our new modelling approach to show that low temperature water-rock reactions are unlikely to have a significant effect on water S18O values. In both cases, the water S18O shift can be explained by oxygen isotope exchange with CO2. Oxygen isotope values observed in the waters measured at Daylesford and Pah Tempe springs are close to equilibrium with S18O of the erupting CO2. Deviation from ideal equilibrium is likely due to localised CO2 movement through the water and the establishment of partial equilibration or kinetic isotope fractionation on degassing bubbles sampled at the water surface.

547 Traditionally, enrichment in 18O in the reservoir waters relative to the MWL has been interpreted to be the

548 result of geothermal activity, while 18O depletion is proposed to be due to CO2-water interaction at lower

549 temperatures. Our global dataset of oxygen and hydrogen isotope measurements in waters from low

550 temperature CO2 springs and the case studies presented from the Daylesford and Pah Tempe CO2 springs

551 provide evidence that equilibration with CO2 can result in both 18O enrichment and depletion in spring waters

552 and therefore geothermal conditions are not necessary to produce 18O-enriched waters. This should be

553 considered in future studies, and used to re-interpret data from previous studies using the water and mineral

554 stable isotope composition to infer water circulation depths, temperatures and local tectonic settings.

555 Acknowledgements

556 R. Karolyte acknowledges the support of an EPSRC PhD studentship in partnership with CO2CRC and

557 Badley Geoscience Ltd. S. Serno was funded by the UK Carbon Capture and Storage Research Centre

558 (UKCCSRC) Call 2 grant, G. Johnson and S. Gilfillan were partially supported by both UKCCSRC and

559 Scottish Carbon Capture and Storage (SCCS). Steve Nelson of Brigham Young University, Utah is thanked

560 for introducing Gilfillan to the Pah Tempe springs and assisting with gas sample collection and background

561 data. We thank Allan Chivas for his assistance in the field during collection of samples for the Daylesford

562 springs and obtaining measurements of water samples. Ian Cartwright is thanked for providing background

563 data on the Daylesford springs. We thank Terry Donnelly, Adrian Boyce and Tony Fallick at SUERC for their

564 assistance in obtaining stable isotope measurements of gas samples.

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Highlights

• Natural S18O deviations in CO2-rich waters are observed around the world.

• New method presented to decouple the influence of water-rock reactions and isotope exchange with CO2 on water S18O variation.

• Low temperature water-rock reactions are unlikely to alter water S18O.

• Oxygen isotope exchange with CO2 can either deplete or enrich natural water in 18O

• Geothermal conditions are not necessarily required to produce 18O-enriched waters.