Stable carbon and oxygen equilibrium isotope fractionation of supercritical and subcritical CO2 with DIC and H2O in saline reservoir fluids

International Journal of Greenhouse Gas Control 39 (2015) 215–224

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Stable carbon and oxygen equilibrium isotope fractionation of supercritical and subcritical CO2 with DIC and H2 O in saline reservoir fluids Veith Becker a,b,∗ , Anssi Myrttinen b , Michael Nightingale a , Maurice Shevalier a , Luc Rock c , Bernhard Mayer a , Johannes A.C. Barth b a

Applied Geochemistry group, Department of Geoscience, University of Calgary, 2500 University Drive NW, Calgary, Alberta T2N 1N4, Canada GeoZentrum Nordbayern, Department of Geography and Earth Sciences, Friedrich-Alexander University Erlangen-Nuremberg (FAU), Schloßgarten 5, 91054 Erlangen, Germany c Shell Canada Limited, 400 4th Ave SW, Calgary, Alberta T2P 2H5, Canada b

a r t i c l e

i n f o

Article history: Received 25 November 2014 Received in revised form 5 May 2015 Accepted 11 May 2015 Available online 1 June 2015 Keywords: Stable isotopes CO2 storage Supercritical CO2 Saline fluids Isotope fractionation

a b s t r a c t The stable isotope composition of CO2 is often used as a tracer during carbon storage projects. To date it has not been investigated to what extent the transition from supercritical to subcritical CO2 affects stable isotope fractionation of CO2 with dissolved inorganic carbon (DIC) and H2 O at elevated temperatures and pressures. We determined the influence of the supercritical state of CO2 on stable carbon and oxygen equilibrium isotope fractionation between CO2 and two types of saline waters: (a) a 80 g/L total dissolved solids (TDS) saline formation water from the Midale formation of the Weyburn CO2 Monitoring and Storage Project (2260 mg/L DIC); (b) a synthetic, DIC free NaCl saline brine with 250 g/L TDS, similar to fluids in the Basal Cambrian Sandstone targeted by the Shell Quest project in Alberta, Canada. The laboratory equilibration experiments between CO2 and saline water were conducted at pressures from 1.0 to 9.0 MPa and temperatures from 22 to 86 ◦ C. We found that oxygen isotope fractionation between CO2 and H2 O (18 OCO2 −H2 O ) for both investigated solutions ranged from 29.0 to 41.1‰ VSMOW and was generally about 1‰ lower than previously reported values for pure water. This discrepancy is likely due to salt effects. Also, 18 OCO2 −H2 O was found to be identical at a given temperature irrespective of whether supercritical or subcritical CO2 was present. Supercritical CO2 did not result in carbon isotope effects that are different from those previously reported between sub-critical CO2 and DIC (13 CDIC−CO2 ). We conclude that supercritical conditions with respect to CO2 in or above CCS storage reservoirs do not cause additional isotope effects and hence do not compromise the use of stable isotopes as a tracer in CO2 storage projects. © 2015 Elsevier Ltd. All rights reserved.

1. Introduction The injection and subsequent long-term storage of CO2 in saline aquifers or depleted oil or gas reservoirs is a suitable approach to mitigate anthropogenic CO2 emissions to the atmosphere and is currently being investigated at various pilot and commercialscale sites worldwide (Michael et al., 2010; Nowak et al., 2013). The long-term feasibility of CO2 storage requires standardized and reliable reservoir and leakage monitoring of injected CO2 at predictable and justifiable cost and effort (Haszeldine, 2009). Various

∗ Corresponding author at: Applied Geochemistry group, Department of Geoscience, University of Calgary, 2500 University Drive NW, Calgary, Alberta T2N 1N4, Canada. Tel.: +1 403 210 8860. E-mail address: [email protected] (V. Becker). http://dx.doi.org/10.1016/j.ijggc.2015.05.020 1750-5836/© 2015 Elsevier Ltd. All rights reserved.

monitoring approaches have been employed at CO2 storage sites worldwide to track the movement and the fate of injected CO2 including geophysical (seismic, geoelectric, electromagnetic techniques), geochemical, and satellite-based approaches (InSAR etc.) among others. For instance, at the 1 Mega Tonnes (MT) CO2 /yr scale Sleipner project in Norway, a number of geophysical monitoring techniques have been used (Chadwick et al., 2006). At the In Salah CO2 project in Algeria (1 MT CO2 /yr), 3D and 4D seismic surveys, geochemical approaches, and satellite-based technologies have been deployed for monitoring of the movement of CO2 in the reservoir (Mathieson et al., 2011; Shi et al., 2012). These approaches are used to monitor subsurface plumes of CO2 but provided only limited insights into geochemical reactions of injected CO2 such as the short- and long-term CO2 –water–rock interactions that may potentially result in solubility and mineral trapping of CO2 in the target reservoir.

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To address these limitations geochemical monitoring based on chemical and stable isotope techniques has been implemented at various CO2 storage field sites and pilot projects to trace the migration and fate of injected CO2 within the reservoir with special focus on CO2 trapping processes (de Caritat et al., 2012; Johnson et al., 2011; Kharaka et al., 2006; Myrttinen et al., 2012b, 2010b; Raistrick et al., 2006). A pre-requisite for this approach is that the isotopic composition of the injected CO2 is distinct from that of baseline (i.e., present before injection) carbon species in the target reservoir as was demonstrated for the Pembina Cardium CO2 pilot in Alberta, Canada (Johnson et al., 2011) and the Weyburn CO2 Monitoring and Storage Project in Saskatchewan, Canada (Emberley et al., 2005; Mayer et al., 2013; Raistrick et al., 2006). A quantitative determination of CO2 movement, trapping mechanisms and reaction pathways based on stable isotope tracer techniques requires accurate knowledge of the carbon and oxygen isotope fractionation for the relevant reactions at conditions occurring in storage reservoirs. These processes have the potential to alter the isotope composition of CO2 in storage reservoirs in systematic ways and thus knowledge of potential isotope fractionation effects is required for the accurate interpretation of isotope monitoring data (Becker et al., 2011). This is true for monitoring of CO2 within the reservoir itself and in overlying strata for verification of containment. The relevance of isotope fractionating geochemical reactions and physical processes that will affect carbon and oxygen stable isotope compositions in CO2 storage scenarios including gas-water equilibration, mineral dissolution and precipitation, sorption, diffusion, and pressure and temperature changes is described and reviewed in Mayer et al. (2015). Conditions in CO2 storage reservoirs may exceed temperatures of 120 ◦ C and pressures of 25 MPa at total dissolved solids concentrations (TDS) of reservoir fluids of up to 250 g/kg (Kühn et al., 2012). At these elevated temperature and pressure conditions, CO2 occurs in a supercritical state. To our knowledge, carbon isotope fractionation factors between CO2 and dissolved inorganic carbon (DIC) and oxygen isotope fractionation between CO2 and H2 O have not been fully verified at pressures and temperatures that encompass the transition from CO2 in subcritical to supercritical state. DIC speciates into H2 CO3 * (defined as the sum of H2 CO3 and CO2(aq) ), HCO3 − and CO3 2− with their relative abundance depending on the pH of the fluid. Depending on the prevailing DIC species, validated and experimentally derived carbon isotope fractionation data between CO2 and DIC are either not available or not fully verified for temperatures above 120 ◦ C for H2 CO3 * (Myrttinen et al., 2014), above 70 ◦ C for HCO3 − (Mook et al., 1974) and above 40 ◦ C for CO3 2− (Myrttinen et al., 2012a). Studies of carbon isotope fractionation between CO2 and individual DIC species have been conducted by other authors yielding fractionation factors for higher temperatures than those listed above, e.g. Malinin et al. (1967) for HCO3 − up to 286 ◦ C and Halas et al. (1997) for CO3 2− up to 200 ◦ C. While the latter study extrapolated isotope fractionation factors from lower temperatures, results of Malinin et al. (1967) were derived from experiments but show an offset to data from other studies (Mook et al., 1974; Szaran, 1997; Zhang et al., 1995) at lower temperatures. Both datasets were thus deemed not validated in the scope of this study. Previous studies of carbon isotope fractionation between DIC and CO2 are described and discussed in detail in Myrttinen et al. (2012a). Oxygen isotope fractionation between CO2 and water is pH independent and its extent is accurately known for pure water up to temperatures of 100 ◦ C (Bottinga, 1968). However, combined pressure and temperature conditions that result in supercritical CO2 (CO2(sc) ) and salinity effects encompassing TDS concentrations and geochemical compositions typical for formation waters in CO2 storage reservoirs have not been fully investigated. It was previously reported that the ionic strength and composition of aqueous solutions may potentially result in oxygen isotope fractionation offsets

higher than analytical precision compared to pure water (Fortier, 1994; Lécuyer et al., 2009; O’Neil and Truesdell, 1991; Sofer and Gat, 1972; Truesdell, 1974). Combinations of high temperatures, elevated solute compositions (salinity), and high pressures (total pressures and pCO2 ) may have effects on equilibrium isotope fractionation, which make extrapolation of isotope fractionation factors from known temperature, pressure and salinity regions questionable (Hoefs, 2009). This includes possible effects of the transition of gaseous CO2 to its supercritical state when ambient conditions exceed the critical point for CO2 at 31.1 ◦ C and 7.39 MPa (e.g. Bachu, 2003). Since CO2 is usually stored in geological reservoirs in supercritical state, it may transition to subcritical state during potential leakage towards the near-surface environment. This phase change must be considered to reliably use stable isotope methods as tracers for CO2 in storage reservoirs that are typically located at depths between 800 m and 3000 m below ground (Kühn et al., 2012). Therefore, the objective of this laboratory study was to determine carbon isotope fractionation effects between CO2 and DIC, and oxygen isotope fractionation effects between CO2 and H2 O in saline reservoir waters at pressure and temperature conditions that cover the transition from supercritical to subcritical CO2 . 2. Methods In order to investigate isotope effects under elevated temperatures and pressures, laboratory experiments are preferred over in-situ investigations because they allow samples to be obtained more frequently and with increased accuracy under controlled conditions. To investigate carbon isotope fractionation between CO2 and DIC and oxygen isotope fractionation between CO2 and reservoir H2 O, a total of 60 equilibration experiments between CO2 and saline waters were conducted in heatable pressure vessels. Two different experimental setups were used: 1) a supercritical fluid extraction reactor (SFER) and 2) a setup composed of gas sampling cylinders. 2.1. Super-critical fluid reactor (SFER) The supercritical fluid extraction reactor (SFER) is a heatable pressure reactor that consists of a 100 mL, high-pressure stainless steel reaction vessel rated for pressures of up to 10,000 psi (∼68.9 MPa). The vessel is placed in an insulated heat chamber, attached to a liquid-CO2 pump via 1/16 high pressure metal tubing. The inlet tubing connects the pump to the top while the outlet tubing connects to the bottom of the vessel via 1/4 NPT (National Pipe Thread) connectors (Fig. 1). The vessel is pressurized via a digitally controlled dual piston CO2 pump. A needle valve between the pump and the vessel allows sealing the vessel from the pump for batch experiments. A digitally controlled heater is installed in the bottom part of the device housing and allows heating the main chamber to temperatures of up to 200 ◦ C. A second heater maintains a constant temperature at the outlet port to prevent cooling due to depressurization during sampling. The sampling port is equipped with two valves, one needle valve to shut off the sampling port and one restrictor valve to adjust the flow rate towards the sampling port. The sampling port ends in a tapered piece of 1/8 tubing that can be connected to a suitable sampling container. 2.2. Sampling cylinders A second experimental setup was developed to complement the experiments with the SFER. Here the objective was to establish a more flexible system that enabled us to conduct multiple paral-

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Table 1 Compositional and stable carbon (ı13 CDIC ) and oxygen (ı18 OH2 O ) isotope data of the original brine fluid from the Midale formation in Weyburn, Saskatchewan, Canada. Formation

Cl mg/L

Br

SO4

Si

Ba

Sr

Li

Na

Ca

Mg

Mn

Fe

K

Midale Location

43747 EC mS/cm 102.9

102 RW m 0.1

3347 pH

17 Alk mg/L 2260

0.5 d13 CDIC ‰ −11.5

64 d18 OH2 O

12 TDS mg/L 79213

26163

2056

514

0.2

1.9

728

Weyburn, SK

6.6

−5.6

Fig. 2. Schematic of the experimental setup using interconnected gas sampling cylinders. In the main reaction vessel with pressure gauge, CO2 /fluid equilibration takes place. The sampling vessel is used to separate fluid and gaseous/supercritical phase before depressurization.

Fig. 1. Schematic of the supercritical fluid extraction reactor (SFER).

lel experiments in addition to the SFER. The main reaction vessel was a SwagelokTM 500 mL PTFE lined stainless steel double-ended sample cylinder (Swagelok Part No.: 304L-HDF4-500-T) equipped with a pressure gauge (Swagelok Part No.: PGI-50M-PG3000-LAOX, 0–3000 PSI ± 2.5%) on one end and a needle valve (Swagelok Part No.: SS-1KM4) on the other. All parts were pressure rated to at least 1800 psi (12.4 MPa). Up to four of these vessels were placed on a shaker and agitated daily for 1 h to ensure homogeneous mixing of the phases in the reaction chamber and thus helped to minimize equilibration times. For rapid phase separation (<2 s) during sampling, an evacuated sampling vessel (Swagelok Part No.: 304HDF4-50) of a volume equal to the amount of fluid used (50 mL, cf. Section 2.3) and equipped with needle valves on both ends was attached to the main reaction vessel via a Hex Coupling (Swagelok Part No.: SS-4-HCG) at the end of the experiments (Fig. 2). Elevated reaction temperatures were maintained with heat tapes that were wrapped around the reaction vessel. Temperatures were measured on the outside of the vessels during experiments. Temperature readings were calibrated beforehand for the temperature difference between inside and outside of the vessel for temperatures from 25 to 90 ◦ C. All temperatures given here are calibrated values for the vessel’s interior. These accurate reaction temperatures thus show minor deviations from the set temperatures of the heat tapes. Maximum experimental temperatures for this setup were 95 ◦ C. Although this setup was developed mostly for subcritical

experiments, the applied p/T ranges also cover the lower range of supercritical conditions for CO2 . 2.3. Experimental procedures The equilibration experiments were conducted in three series, A–C, with fluid volumes of 50 mL each. In series A of the experiments, de-ionized water was used to conduct control tests on the setup and sampling procedures as well as to determine oxygen isotope equilibration times between CO2 and H2 O, since kinetics of isotope equilibration are known to be slower than those for nonisotopic, i.e., elemental equilibrium (Mills and Urey, 1940). Carbon isotope equilibration between CO2 and DIC is sufficiently fast to be considered essentially instantaneous on the time scales relevant to carbon storage monitoring (Becker et al., 2011) and is known to be established on time scales of hours rather than days in similar experimental setups (Myrttinen et al., 2014). Subsequently, series B was conducted using saline fluids from the Midale formation (TDS = 80 g/L, DIC = 2260 mg/L, Table 1) of the Weyburn CO2 Monitoring and Storage Project (Wilson and Monea, 2004). The DIC baseline concentration of the Midale fluids is not expected to affect the experiments of series A, as equilibrium isotope fractionation (13 CDIC−CO2 ) is independent from baseline isotope ratios. Due to the amount of CO2 in the experiments being orders of magnitudes

Table 2 Parameters of oxygen isotope experiments conducted as part of series A. ‘Baseline’ isotope ratios indicate initial composition of phases, non-baseline ı-values are postexperimental results. ‘Time’ denotes equilibration time. Fluid Type De-ionized H2 O

Time days

◦

T C

P MPa

ı18 OCO2 ı18 OH2 O Baseline in ‰ VSMOW

1 1 2 2 4 4 6 14

22 22 22 22 22 22 22 22

1.1 5.1 1 5 1.1 5 1 6

21.6 21.6 21.6 21.6 21.6 21.6 21.6 23

−19.4 −19.4 −19.4 −19.4 −19.4 −19.4 −19.4 −8.8

ı18 OCO2 In ‰ VSMOW

ı18 OH2 O

ı18 OCO2 −H2 O

20.6 21.2 21 21.5 21.5 22 22.2 30.5

−19.3 −18.8 −19.3 −18.6 −19.2 −18.5 −18.7 −10.2

39.9 40 40.2 40.2 40.6 40.3 40.9 40.7

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Fig. 3. Oxygen isotope fractionation between subcritical CO2 and de-ionized H2 O from experiments in sample cylinders and the SFER together with equilibrium fractionation data from Bottinga (1968). The reaction temperature was 22 ◦ C; Sampling cylinders: 50 mL H2 O was equilibrated with 450 mL CO2 at pCO2 of 1 and 5 MPa; baselines were ı18 OH2 O = −19.4‰ and ı18 OCO2 = 21.6‰ vs MOW. SFER: 50 mL H2 O was equilibrated with 50 mL CO2 at pCO2 = 6.0 MPa; baselines were ı18 OH2 O = −8.8‰ and ı18 OCO2 = 23.0‰ vs MOW.

larger than the amount of DIC, its influence on ı13 CCO2 and ı13 CDIC will be insignificant for the experiments presented here. For series C, an artificial 250 g/L NaCl solution was prepared with de-ionized water that contained no DIC. The salinity of this solution is similar to that of formation waters in the Basal Cambrian Sandstone, the target zone for CO2 injection at the Shell Quest saline aquifer CO2 injection project in Alberta, Canada. Vessels were flushed with deionized water between all experiments and washed with CH2 Cl2 (dichloromethane) and HNO3 after experiments with saline solutions in order to remove any salt and organic residues. The vented and dried vessels were filled with 50 mL of the relevant liquid for all experiments in series A–C and subsequently sealed and pressurized with CO2 (99.9% purity). The applied pCO2 at experimental temperatures ranged from 1.0 to 9.0 MPa. The vessels were sealed and pressurized prior to heating. The pressurization procedure was kept short (within less than 3 min) rather than to meet exact target pressures to prevent isotope exchange between water and gases in the vessel and external reservoirs (gas tanks and tubing). This approach and expansion during temperature increase lead to reaction pressures that varied slightly around a target value. Baseline isotopic data for the experiments are summarized in Tables 2 (series A), 3 (oxygen isotope data for series B), 4 (oxygen isotope data for series C) and 5 (carbon isotope data for series B), respectively. Three different tanks of CO2 were used with ı13 C values of −37.9, −30.6 and −3.3‰ VPDB and ı18 O values of +21.6, +23.0 and +26.3‰ VSMOW, respectively. DIC in the Midale saline formation water had ı13 CDIC values of −9.1, −9.7 to −11.5‰ VPDB before the experiments started (i.e., baseline, series B). The deionized water used for series A and the NaCl solution used in series C were DIC free. ı18 OH2 O values of deionized water used in in series A were −19.4 and −8.8‰ VSMOW, respectively. The Midale formation water used in series B had a ı18 OH2 O value of −5.7‰ VSMOW and the ı18 OH2 O value of the NaCl solution used in series C was −19.4‰ VSMOW. For series A experiments, conditions were chosen, for which literature data for carbon isotope fractionation between CO2 and DIC (Myrttinen et al., 2012a) and oxygen isotope fractionation between CO2 and H2 O (Bottinga, 1968) was available. These experiments were conducted at a temperature of 22 ◦ C and pressures of 1 and 5 MPa with run times from 1 to 14 days. Based on the results of series A, showing equilibrium values for oxygen isotope fractionation after more than 5 days (Fig. 3 and Table 2), runtimes of at least 6 days were chosen for experiments of series B and C to ensure achievement of isotopic equilibrium. Temperatures and pressures

in series A ranged from 22 to 86 ◦ C and 1.0 to 8.8 MPa for series B and from 22 to 82 ◦ C and 2.0 to 9.0 MPa for series C. Carbon and oxygen isotope effects were investigated in separate parallel experiments of series B. This was required to adapt sampling strategies for each isotope system as described below. Subsampling of fluid samples from a single experiment for analysis of both carbon and oxygen isotope ratios would have facilitated degassing of CO2 from the sample and subsequent partitioning of carbon isotopes between DIC and CO2 would have potentially affected the ı13 CDIC . At the end of the experiments, CO2 gas samples were collected into one-litre PVF (polyvinyl fluoride) bags for subsequent isotope analysis. For ı13 CDIC analysis, the total fluid volume was transferred from the sampling vessel directly and quantitatively into a separate one-litre PVF bag to avoid contamination from ambient air. Degassing CO2 from the aqueous samples within the sampling bag did not affect ı13 CCO2 values as CO2 was precipitated together with the DIC as SrCO3 within the sampling bag prior to analysis as described in Section 2.4. For analysis of ı18 O values of water, 2 mL of fluid was transferred to glass vials from the experiments. Degassing CO2 from depressurized samples had no effect on ı18 OH2 O values due to the negligible amount of oxygen in degassing CO2 compared to H2 O. This was verified by mass balance calculations. Sampling of the SFER was conducted at experimental temperatures. Due to the single-vessel design of the SFER, sampling resulted in depressurization of the sampling vessel and thus was conducted within seconds to avoid re-equilibration at reduced pressure. After the H2 O phase was sampled, the remaining H2 O was drained from the vessel. The CO2 phase was sampled immediately afterwards. Subsequent sampling of the remaining CO2 phase into different gasbags confirmed that no isotope fractionation within the CO2 phase occurred during degassing of the vessel. In the sampling cylinder setup, the phases were separated as described in Section 2.2 immediately after removal of the heat tapes and before vessel temperatures decreased significantly. Reequilibration of the phases was thus precluded. 2.4. Analytical methods Alkalinity and pH of the water samples were measured using a Thermo Scientific® Orion 940 auto-titrator. Fluid samples for ı13 CDIC analysis were treated with ammoniacal SrCl2 solution to quantitatively precipitate DIC as SrCO3 . SrCO3 was vacuum filtered, washed and dried and subsequently converted to CO2 with 105% H3 PO4 and analysed for ı13 CDIC on a dual inlet isotope ratio mass spectrometer (IRMS) built with Micromass 903 components. ı13 CCO2 and ı18 OCO2 values were measured on a ThermoFisher MAT 253 IRMS coupled to GC-C (Isolink, ThermoFisher). Water samples were analysed for ı18 OH2 O on a ThermoFisher Delta V plus IRMS coupled to a Thermal Combustion Elemental Analyzer (TC/EA) via pyrolysis that produced CO. The measured stable isotope ratios are expressed as permille deviations from the international standards Vienna Pee Dee Belemnite (VPDB) for carbon isotopes and Vienna Standard Mean Ocean Water (VSMOW) for oxygen isotopes expressed by the delta notation as follows (Clark and Fritz, 1997): ı(‰) = ((H/L sample – H/L standard )/H/L standard ) × 1000

(1)

13 C, 18 O)

and light (i.e. where H and L represent the heavy (i.e. stable isotopes of carbon and oxygen, respectively. External reproducibility, i.e. the standard deviation (1 ) of lab-internal standards was <±0.5‰ for ı13 CCO2 and ı18 OCO2 and <±0.2‰ for ı18 OCO2 and ı13 CDIC . ı18 OCO2 results were converted from VPDB to VSMOW scale according to Coplen et al. (1983): 12 C, 16 O)

ı18 OVSMOW = 1.03091 × ı18 OVPDB + 30.91(‰)

(2)

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Table 3 Parameters of oxygen isotope experiments conducted as part of series B. ‘Baseline’ isotope ratios indicate initial composition of phases, non-baseline ı-values are postexperimental results. ‘State’ indicates if experimental conditions were sub- or supercritical with respect to CO2 . Fluid type 80 g/L Midale formation fluid

T C

P MPa

State

◦

ı18 OCO2 Baseline in ‰ VSMOW

ı18 OH2 O

ı18 OCO2 In ‰ VSMOW

ı18 OH2 O

18 OCO2 −H2 O

22 22 22 22 22 22 22 22 51 52 56 57 59 62 70 57 58 59 71 86

1 1.1 1.1 1.1 1.1 4 4 4.9 7.1 1.1 3.9 5.4 3.8 1.1 6 8 8 7.6 8.1 8.8

sub sub sub sub sub sub sub sub sub sub sub sub sub sub sub sc sc sc sc sc

21.6 21.6 21.6 26.3 26.3 26.3 26.3 21.6 26.3 26.3 26.3 21.6 26.3 26.3 21.6 21.6 21.6 26.3 26.3 26.3

−5.7 −5.7 −5.7 −5.7 −5.7 −5.7 −5.7 −5.7 −5.7 −5.7 −5.7 −5.7 −5.7 −5.7 −5.7 −5.7 −5.7 −5.7 −5.7 −5.7

32.7 33.3 31.1 32.4 31.6 30.4 30.5 30.2 28 27.9 26.6 23.3 26.4 26.5 22.7 24.2 24.5 27.5 25.7 26

−7.6 −7.8 −7.8 −7.2 −7.2 −8.7 −8.9 −11.2 −6.5 −6.5 −6.3 −9.8 −6.1 −7 −8.7 −9 −9.2 −5.7 −4.8 −3.9

40.3 41.1 39 39.6 38.8 39.1 39.4 39.9 34.4 34.5 32.9 33.1 32.4 33.6 31.4 33.2 33.7 33.1 30.5 29.9

Isotope fractionation is caused by slightly different equilibrium constants for the different isotopes involved in a specific physico-chemical reaction. These changes in stable isotope ratios are commonly denoted by the isotope fractionation factor ˛, which is defined as ˛X−Y =

  ıX + 103   3

(3)

ıY + 10

for the phases or compounds X and Y between which isotope fractionation occurs. The extent of isotope equilibrium fractionation between two phases or compounds is independent of baseline isotope ratios or concentrations and is often reported as 103 ln˛, approximating the isotopic difference  (also called the isotope enrichment factor) between the ı values of two reacting compounds in ‰. 3. Results 3.1. Oxygen isotopes Experiments of series A were conducted at 22 ◦ C with runtimes of 1, 2, 4, 6, and 14 days, and results are summarized in Table 2. The 7 experiments with runtimes of up to 6 days were conducted in the sampling cylinder setups, while the 14 day experiment was conducted in the SFER. All results are expressed on the VSMOW scale and ı18 OH2 O values of de-ionized water prior to reaction in 7

of 8 experiments were -19.4‰, while the CO2 had a ı18 OCO2 value of +21.6‰ for experiments using sampling cylinders. After reaction for 1–6 days at pressures between 1.0 and 5.1 MPa, the released CO2 had ı18 OCO2 values from +20.6 to +22.2‰ while the ı18 OH2 O values ranged between −19.3 and −18.5‰ at the end of the experiments. Hence, the ı18 OCO2 values were between +39.9 and +40.9‰ higher than those of H2 O. The one experiment conducted in the SFER contained water with a baseline ı18 OH2 O value of −8.8‰ and a ı18 OCO2 value of +23.0‰. After 14 days of reaction at 6.0 MPa, the ı18 O value of the released CO2 had increased to +30.5‰ while the water ı18 O value had decreased to −10.2‰. The difference between ı18 O values of CO2 and H2 O at the end of this experiment (18 OCO2 −H2 O ) was thus +40.7‰. Experiments of series B using Midale formation water were conducted at temperatures from 22 to 86 ◦ C at pCO2 from 1.0 to 8.8 MPa for at least 6 days, and results are summarized in Table 3. The ı18 OH2 O value of saline Midale formation water prior to reaction was −5.6‰. Two different CO2 sources were used with ı18 OCO2 values of +21.6 and +26.3‰, respectively. Subcritical equilibration experiments of series B covered temperatures from 22 to 70 ◦ C at pCO2 from 1.0 to 7.1 MPa. At the end of the experiment, ı18 OCO2 values ranged from +22.7‰ at 70 ◦ C to +33.3‰ at 22 ◦ C and ı18 OH2 O values varied from -11.2‰ at 22 ◦ C to -6.1‰ at 59 ◦ C. Supercritical experiments covered temperatures from 57 to 86 ◦ C at pCO2 from 7.6 to 8.8 MPa. At the end of these

Table 4 Parameters of oxygen isotope experiments conducted as part of series C. ‘Baseline’ isotope ratios indicate initial composition of phases, non-baseline ı-values are postexperimental results. ‘State’ indicates if experimental conditions were sub- or supercritical with respect to CO2 . Fluid Type 250 g/L NaCl-solution

T C

P MPa

State

◦

ı18 OCO2 ı18 OH2 O Baseline in ‰ VSMOW

22 22 42 44 44 45 45 46 48 48 82 76

3 5.9 3.2 2 2.1 4.3 7 7 5 6.3 4.1 9

sub sub sub sub sub sub sub sub sub sub sub sc

26.3 26.3 26.3 26.3 26.3 26.3 26.3 26.3 26.3 26.3 26.3 26.3

−19.4 −19.4 −19.4 −19.4 −19.4 −19.4 −19.4 −19.4 −19.4 −19.4 −19.4 −19.4

ı18 OCO2 In ‰ VSMOW

ı18 OH2 O

18 OCO2 −H2 O

23.3 24.5 19.9 19.3 20.2 20.8 22.9 23.2 21 22.6 16.1 19.8

−16.4 −15.1 −15.4 −16.3 −16.2 −14.3 −13.1 −12.5 −13.9 −12.7 −13 −10.5

39.6 39.7 35.3 35.7 36.3 35.1 36 35.8 35 35.3 29 31.3

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Table 5 Parameters of carbon isotope experiments conducted as part of series B. ‘Baseline’ isotope ratios indicate initial composition of phases, non-baseline ı-values are postexperimental results. ‘State’ indicates if experimental conditions were sub- or supercritical with respect to CO2. ‘pH’ is reconstructed from fractionation data after Myrttinen et al. (2015). Values for  are calculated based on non-rounded delta values. Fluid Type 80 g/L Midale formation fluid

T C

P MPa

State

ıı13 CCO2 ı13 CDIC Baseline in ‰ VPDB

ı13 CCO2 In ‰ VPDB

ıı13 CDIC

13 CDIC−CO2

pH

◦

22 22 22 22 22 22 22 22 55 56 59 61 71 72 77 68 71 76 78 86

1 1 1.1 1.4 4.1 4.8 6 6 3.5 3.9 3.8 3.9 7 7 7.2 8 8.1 8.4 7.4 8.8

sub sub sub sub sub sub sub sub sub sub sub sub sub sub sub sc sc sc sc sc

−37.9 −37.9 −37.9 −37.9 −3.3 −37.9 −37.9 −37.9 −3.3 −3.3 −3.3 −3.3 −37.9 −37.9 −37.9 −3.3 −3.3 −3.3 −37.9 −3.3

−38.5 −37.9 −38.3 −38 −3.7 −37.8 −38 −37.8 −3.5 −3.5 −3.5 −3.6 −38.3 −38.3 −38.3 −3.6 −3.6 −3.6 −38.3 −3.6

−34.4 −34 −34.4 −33.8 −3 −35.5 −38 −38 −3.7 −3.4 −3.8 −4.1 −38.2 −37.6 −37.1 −4.3 −4.2 −4.2 −37.6 −4.1

4 3.8 3.9 4.2 0.8 2.2 0 −0.2 −0.2 0.1 −0.2 −0.5 0.1 0.6 1.2 −0.7 −0.6 −0.6 0.7 −0.5

6.5 6.4 6.4 6.5 5.8 6.1 5.5 5.4 5.4 5.6 5.4 5.3 5.7 6 6.3 5 5.1 5.3 6.1 5.4

experiments ı18 OH2 O ranged from +24.2‰ at 57 ◦ C to +27.5‰ at 59 ◦ C and ı18 OH2 O values varied from −9.2‰ at 58 ◦ C to −3.9‰ at 86 ◦ C. Oxygen isotope enrichment factors (18 OCO2 −H2 O ) for series B ranged from +31.4‰ at 70 ◦ C to +41.1‰ at 22 ◦ C for experiments with subcritical CO2 and from +29.9‰ at 86 ◦ C to +33.7‰ at 58 ◦ C for experiments with CO2(sc) . Both supercritical and subcritical experiments thus showed a general trend to lower 18 OCO2 −H2 O values at higher temperatures. Experiments of series C using 250 g/L NaCl solution were conducted at temperatures from 22 to 82 ◦ C and at pCO2 values ranging from 2.0 to 9.0 MPa for at least 6 days, and results are summarized in Table 4. The ı18 OH2 O value of the artificial brine prior to reaction was −19.4‰. The initial ı18 OCO2 of the CO2 was +26.3‰. Experiments of series C under subcritical conditions with respect to CO2 covered temperatures from 22 to 82 ◦ C at pCO2 from 2.0 to 7.0 MPa. After equilibration, ı18 OCO2 values ranged from +16.1‰ at 82 ◦ C to +24.5‰ at 22 ◦ C and ı18 OH2 O values varied from 16.4‰ at 22 ◦ C to -12.5‰ at 46 ◦ C. One supercritical experiment was conducted at 76 ◦ C and 9.0 MPa. Here, the resulting ı18 OCO2 after equilibration was +19.8‰ and ı18 OH2 O was -10.5‰. Oxygen isotope enrichment factors (18 OCO2 −H2 O ) for series C ranged from +29.0‰ at 82 ◦ C to +39.7‰ at 22 ◦ C for subcritical experiments and +31.3‰ for the supercritical experiment at 76 ◦ C. Hence, the results of the experiments of series C also displayed a trend to lower 18 OCO2 −H2 O values at higher temperatures. 3.2. Carbon isotopes Stable carbon isotope ratios of CO2 and DIC were measured for the experiments with Midale saline waters (series B) and are listed in Table 5. DIC in the Midale saline formation water at the beginning of the experiments had ı13 CDIC values of −9.1, −9.7 and −11.5‰. Two different CO2 sources were used with ı13 CCO2 values of −37.9 and −3.3‰, respectively. This large difference in baseline ı13 CCO2 values has a major impact on the resulting carbon isotope ratios of DIC and CO2 . Therefore the results of the experiments are divided into two groups based on the ı13 CCO2 baseline values. Table 5 shows the results of all experiments of series C sorted by state conditions for CO2 , temperature and pressure. Experiments with a baseline ı13 CCO2 value of −37.9‰ represented predominantly subcritical conditions with temperatures from 22 to 78 ◦ C and pCO2 from 1.0 to 7.4 MPa. At the end of the

−11.5 −11.5 −11.5 −9.1 −9.7 −11.5 −9.1 −9.1 −9.7 −9.7 −9.7 −9.7 −11.5 −11.5 −11.5 −9.7 −9.7 −9.7 −11.5 −9.7

experiments, ı13 C values of CO2 ranged from −38.5 to −37.8‰, while the ı13 C values of DIC ranged from −38.2 to −33.8‰. Experiments with a baseline ı13 CCO2 value of −3.3‰ represented both subcritical and supercritical conditions. With the exception of the experiment run at 22 ◦ C, temperatures ranged from 55 to 86 ◦ C and pCO2 from 3.5 to 8.8 MPa. At the end of the experiments, the ı13 CCO2 values were constant at −3.6 ± 0.1‰ whereas the ı13 CDIC values varied from −4.3 to −3.0‰. Carbon isotope enrichment factors (13 CDIC−CO2 ) over all subcritical experiments of series C ranged from −0.5‰ at 61 ◦ C to +4.2‰ at 22 ◦ C, while supercritical experiments showed a range from −0.5‰ at 68 ◦ C to +0.7‰ at 78 ◦ C. 4. Discussion 4.1. Oxygen isotopes 4.1.1. Oxygen isotope equilibration times Oxygen isotope fractionation between CO2 and H2 O of series A was evaluated to determine the minimum time required to achieve oxygen isotope equilibrium. As shown in Fig. 3, the oxygen isotope enrichment factor between CO2 and H2 O (18 OCO2 −H2 O ) was +39.9‰ after one day and increased from +40.2‰ after 2 days and +40.5‰ after four days asymptotically to a constant value of +40.8 ± 0.1‰ for experiments with equilibration times of 6 days or more. This value is consistent with published oxygen isotope equilibrium fractionation data between CO2 and H2 O at 22 ◦ C of +40.8‰ (Bottinga, 1968). Therefore, we conclude that oxygen isotope equilibrium was achieved after six days using the applied experimental setups. Even after experimental runtimes of 4 days, the results fall within one standard deviation of the analytical error (0.2‰) of the equilibrium value. These equilibration times for oxygen isotopes between CO2 and H2 O are consistent with findings from Johnson and Mayer (2011), who used a similar setup and found a reaction time of 7 days sufficient to establish stable oxygen isotope equilibrium between CO2 and H2 O. Oxygen isotope equilibrium in series A was achieved for both the SFER and the sample cylinder setups with equilibration times of at least 6 days. Although for both setups, H2 O and CO2 with different baseline values for ı13 C have been used, the resulting 18 OCO2 −H2 O values were identical (within 1 ␴ analytical error) with +40.9‰ for the sampling cylinders and +40.7‰ for the SFER at reaction

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Fig. 4. Oxygen isotope fractionation between subcritical and supercritical CO2 in equilibrium with 80 g/L Midale fluid and 250 g/L NaCl solution. Calculated fractionation for pure water (Bottinga, 1968) for comparison. Regression includes both subcritical and supercritical data. pCO2 between 1.0 and 9.0 MPa. Errorbars show analytical error of 0.5‰ (1 SD).

temperatures of 22 ◦ C. The excellent agreement of the experimentally determined equilibrium oxygen isotope fractionation factors between de-ionized water and CO2 with those reported in the literature validates the experimental setups and sampling procedures at temperature and pressure conditions, for which oxygen isotope fractionation data are available. This provides the foundation to investigate carbon and oxygen isotope fractionation at conditions, for which no validated isotope fractionation factors between CO2 , DIC and saline reservoir waters are available. 4.1.2. Effect of water salinity on oxygen isotope fractionation Fig. 4 displays the observed oxygen isotope enrichment factors between CO2 and H2 O (18 OCO2 −H2 O ) for all experiments conducted with 80 g/L TDS Midale saline water and 250 g/L NaCl solution at temperatures between 22 ◦ C and 86 ◦ C and pressures ranging from 1.0 to 9.0 MPa. For series B (Midale fluid), the highest 18 OCO2 −H2 O value of +41.1‰ was observed at the lowest temperature of 22 ◦ C and decreasing 18 OCO2 −H2 O values of as low as +29.9‰ were determined at the highest temperatures of 86 ◦ C, yielding the following polynomial regression: OCO2 −H2 O

18 OCO2 −H2 O [‰] = −4.0e − 5T2 − 0.166T + 43.200withR 2 = 0.95 (4) Results for the experiments of series C, conducted with 250 g/L NaCl solution, showed a very similar trend with the highest 18 OCO2 −H2 O value of +39.7‰ observed again at the lowest temperature of 22 ◦ C. With increasing temperature, 18 OCO2 −H2 O values generally decreased resulting in the lowest value of +29.0‰ at the highest temperature of 82 ◦ C in series C. The temperature dependence of 18 OCO2 −H2 O for series C is described by: OCO2 −H2 O

18 OCO2 −H2 O [‰] = −7.0e − 5T2 − 0.1836T + 43.697withR 2 = 0.99 (5) resulting in very similar trends for series B and C over the investigated temperature range between 20 and 90 ◦ C. The 18 OCO2 −H2 O values from our experiments with saline Midale water and a synthetic NaCl brine plot parallel to values predicted by Bottinga (1968) for pure water but display a rather constant offset of approximately −1.0‰. Solute composition and salinity (TDS) are known to affect oxygen isotope fractionation (O’Neil and Truesdell, 1991). The extent of the salt effect generally

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depends on temperature and can show different and even opposing trends for different ions in solution (e.g. Truesdell, 1974). For temperatures between 25 and 75 ◦ C, Truesdell (1974) found oxygen isotope fractionation between CO2 and a 4 molar NaCl solution to be reduced by approximately 1‰ compared to oxygen isotope fractionation of CO2 with pure water. This is in excellent agreement with the results of series C in this study compared to equilibrium values of 18 OCO2 −H2 O of series A and those from Bottinga (1968). Results of series B yielded a very similar −1‰ offset for 18 OCO2 −H2 O compared to oxygen isotope fractionation between CO2 and H2 O for pure water (Bottinga, 1968) despite lower TDS (80 g/L compared to 250 g/L). As the Midale waters used in series B contain a variety of different ions (Table 1), this is consistent with results reported by Truesdell (1974) for e.g. MgCl2 and MgSO4 solutions showing a more pronounced salt effect than NaCl solutions. The hydration regions surrounding these ions consist of structured, closely bound zones of water enriched in heavier isotope 18 O and less structured, loosely bound zones that are enriched in the lighter isotope 16 O. With the extent of these zones depending on the ion, the ionic strength of the solution, and temperature, this model explains varying salt effects for different solution compositions and temperatures (Truesdell, 1974). Hence, it is not surprising that seemingly contradictory results for the salt effects on 18 OCO2 −H2 O compared to oxygen isotope fractionation between CO2 and H2 O for pure water have been reported. Horita et al. (1995) found no significant influence of NaCl concentrations on oxygen isotope fractionation between liquid and vapour phases of water below 200 ◦ C in experiments with up to 5 molar NaCl solutions at 100 MPa. The salt effect here is also based on the influence of hydration regions around dissolved ions as described above and thus should show similar trends. Lécuyer et al. (2009), however, reported a shift of 18 OCO2 −H2 O to slightly higher values (∼0.3‰ at 80 g/L, ∼0.5‰ at 250 g/L) for increasing concentrations of KCl and sea salt solutions at ambient pressures and temperatures of 40 ◦ C. It is therefore plausible that oxygen isotope fractionation effects caused by higher salinities also depend on the type of salt rather than the salinity alone. Our results summarized in Fig. 4 reveal a decrease of 18 OCO2 −H2 O of ∼−1‰ compared to oxygen isotope fractionation between CO2 and H2 O for pure water (Bottinga, 1968) for both Midale formation water (series B) and a synthetic NaCl solution with 250 g/L TDS (series C) over the investigated temperature range from 22 to 86 ◦ C. In agreement with result from Truesdell (1974) we suggest that a combination of elevated ionic strengths but different solute compositions explains our experimental data. The oxygen isotope fractionation offset of about 1‰ compared to pure water found in our study is in agreement with the maximum offset for 4 m NaCl solutions found by Truesdell (1974). The ∼1‰ offset of the oxygen isotope fractionation in our data, however, is not temperature dependent between 22 and 88 ◦ C, but such a trend might potentially be obfuscated by the scatter of the data. Elevated pressures of 1 and 5 MPa did not affect 18 OCO2 −H2 O in series A suggesting that pressure variations are not the cause for the observed 18 OCO2 −H2 O compared to oxygen isotope fractionation between CO2 and H2 O for pure water. 4.1.3. Effect of CO2(sc) on oxygen isotope fractionation In Fig. 4, oxygen isotope enrichment factors between CO2 and H2 O (18 OCO2 −H2 O ) determined for experiments under supercritical conditions for CO2 at elevated temperatures and pressures are shown in full symbols, whereas results for experiments conducted at sub-critical conditions usually at lower temperatures and pressures are shown in open symbols. Oxygen isotope enrichment factors 18 OCO2 −H2 O for both supercritical and subcritical experiments plot along the same regression described by Eq. (3) without any noticeable offset. Therefore, we conclude that the oxygen isotope enrichment factor between CO2 and H2 O under the experi-

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mental conditions is at a given temperature identical irrespective of whether supercritical or subcritical CO2 is present. This finding is of key importance for the application of stable oxygen isotope methods for monitoring of CO2 storage reservoirs and most notably in leakage monitoring, where CO2 potentially migrates upwards and will eventually transition from supercritical to subcritical state. Our results suggest that this phase transition will not cause any unpredictable effects on the oxygen isotope ratios of CO2 , thus allowing to apply the same isotope fractionation relations for migration and dissolution of CO2 under supercritical and subcritical conditions. 4.2. Carbon isotope fractionation The extent of stable carbon isotope fractionation between DIC in the Midale saline formation waters (series B) and CO2 depends on the speciation of the DIC (see Section 1). As the DIC speciation depends on the pH of the solution, the latter must be precisely known to determine the distribution of DIC species, into which carbon isotopes of CO2 are partitioned (Clark and Fritz, 1997). Ex-situ measurements of the pH of CO2 pressurized fluids are unreliable, since depressurization, e.g. upon opening the reaction vessels, results in degassing of CO2 from the fluid and thus leads to increasing pH values via the carbonate equilibrium (Myrttinen et al., 2010a). Therefore, a recently published method was applied to determine the in-situ pH of fluids in equilibrium with high pCO2 values from stable carbon isotope data (Myrttinen et al., 2015). This method is based on the distinct carbon isotope fractionation between CO2 and the DIC species H2 CO3 * , HCO3 − and CO3 2− and the pH dependence of the DIC speciation. With accurate equilibrium carbon isotope fractionation factors being known from analysis of CO2 and DIC samples, the in-situ pH value of the solution can be reconstructed. The method is described in detail in Myrttinen et al. (2015). The pH values determined with this technique that are reported in Table 5 were used to derive the DIC species distribution in the experimental fluids. This enabled us to compare the carbon isotope enrichment factors between the total DIC (in this case predominantly H2 CO3 * + HCO3 − ) and supercritical or subcritical CO2 (13 CDIC−CO2 ) for individual experiments despite their varying DIC speciation caused by the different temperatures and CO2 pressures in the 60 different experiments. This classification by pH is crucial for comparing isotope fractionation under supercritical and subcritical conditions with respect to CO2 because experimental conditions for sub and supercritical experiments have to be distinct in either temperature or pressure which will always cause slightly differing pH values in corresponding samples that can lead to significant differences in 13 CDIC−CO2 . Fig. 5 displays as dotted and dashed lines the pH dependant carbon isotope fractionation between DIC and CO2 as enrichment factor (13 CDIC−CO2 ) values for temperatures of 20, 60, 80 and 100 ◦ C based on data from Mook et al. (1974) and Myrttinen et al. (2014). In our experiments, highest 13 CDIC−CO2 values of +4.2‰ were observed for experiments conducted at 22 ◦ C and low pressures of 1 MPa. Under these conditions, the fluids in the reaction vessel have a pH value of 6.4 and our results are in very good agreement with literature data at 20 ◦ C. Experiments conducted at higher pressures and temperatures between 22 and 80 ◦ C, while remaining at subcritical conditions for CO2 , resulted in pH values of the saline water between 5.3 and 6.3. Here 13 CDIC−CO2 values between +0.5 and +2.0‰ were observed (Fig. 5) with the higher 13 CDIC−CO2 values associated with higher pH values. Experiments conducted at supercritical conditions for CO2 resulted in pH values ranging from 5.0 to 6.1. Under these conditions 13 CDIC−CO2 values were within 0 ± 1‰. The excellent agreement between results from subcritical and supercritical experiments suggests that there is only minor carbon isotope fractionation between supercritical CO2 and DIC. This is

Fig. 5. Carbon isotope fractionation between sub- and super-critical CO2 and DIC in Midale saline water with reconstructed pH. Lines show calculated, expected carbon isotope fractionation for temperatures of 20, 60, 80 and 100 ◦ C (after Myrttinen et al., 2015). pCO2 between 1.1 and 8.0 MPa. Denoted pressures refer to experiments at 22 ◦ C. Error bars show combined analytical error of 1 SD = 0.5‰.

consistent with pH dependant DIC−CO2 trends predicted based on work by Mook et al. (1974), Vogel et al. (1970) and Myrttinen et al. (2014). Therefore we conclude that the occurrence of supercritical CO2 in our experimental setup did not result in carbon isotope effects that are different from those between subcritical CO2 and DIC. Since pH values of the solutions are typically low at conditions where supercritical CO2 occurs and H2 CO3 * is usually the dominant DIC species, our experiments reveal that carbon isotope fractionation between DIC and CO2 (13 CDIC−CO2 ) is expected to be small (<1‰). This is supported by carbon isotope measurements of series C where ı13 C values of CO2 at the end of the experiments deviated by less than ±0.5‰ from those of the initial CO2 . The initial DIC concentration of the Midale fluids does not directly affect the pH-13 CDIC−CO2 relationship. However, DIC baseline concentrations may affect final pH values due to pH-buffering and thus impact the resulting 13 CDIC−CO2 indirectly. Due to the low amount of DIC compared to the large amount of CO2 this effect is however minimal in our series B experiments. 5. Conclusions The supercritical fluid extraction reactor (SFER) and the sampling cylinders were suitable for conducting stable carbon and oxygen isotope equilibration experiments with saline fluids and CO2 under subcritical and supercritical conditions with respect to CO2 for pressures up to 9.0 MPa and temperatures up to 86 ◦ C. Using both experimental setups we demonstrated that oxygen isotope fractionation between CO2 and H2 O (18 OCO2 −H2 O ) in saline (80 g/L TDS) Midale formation water and a synthetic NaCl brine with 250 g/L TDS was generally around 1‰ lower than previously reported for pure water (Bottinga, 1968). This appears to be due to a salt effect that is influenced by total salinity as well as the solute composition of the saline waters. Regardless of this minor offset, the oxygen isotope enrichment factors between CO2 and H2 O (18 OCO2 −H2 O ) was found to be highly predictable and identical at a given temperature irrespective of whether supercritical or subcritical CO2 was present. Furthermore, the occurrence of supercritical CO2 did not result in carbon isotope effects that are different from those between subcritical CO2 and DIC (13 CDIC−CO2 ). Experiments conducted at supercritical conditions for CO2 resulted in pH values ranging from 5.0 to 6.1 while carbon isotope enrichment factors between DIC and

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CO2 (13 CDIC−CO2 ) were within 0 ± 1‰. This small isotope effect suggests that there is only minor carbon isotope fractionation between supercritical CO2 and DIC. Therefore, we conclude that the occurrence of supercritical CO2 or the transition from supercritical to subcritical CO2 in or above CO2 storage reservoirs does not cause additional carbon and oxygen isotope effects and hence does not compromise the use of stable isotopes to trace the movement and fate of CO2 in geological storage projects. Acknowledgements We gratefully acknowledge funding of this study by the GEOTECHNOLOGIEN Program (Project CO2 ISO-LABEL, Grant No: 03G0801A) of the German Federal Ministry of Education and Research (BMBF); the Quest Carbon Capture and Storage project; and Carbon Management Canada Inc. (CMC-NCE) as a part of the project Storage Geochemistry (C01). The Quest project is a joint venture between Shell Canada Energy, Chevron Canada Limited, and Marathon Oil Canada Corporation, and is operated by Shell. Funding for the Quest CCS project from the Government of Alberta and the Government of Canada is gratefully acknowledged. References Bachu, S., 2003. Screening and ranking of sedimentary basins for sequestration of CO2 in geological media in response to climate change. Environ. Geol. 44, 277–289, http://dx.doi.org/10.1007/s00254-003-0762-9 Becker, V., Myrttinen, A., Blum, P., van Geldern, R., Barth, J.A.C., 2011. Predicting ı13 CDIC dynamics in CCS: a scheme based on a review of inorganic carbon chemistry under elevated pressures and temperatures. Int. J. Greenh. Gas Control 5, 1250–1258, http://dx.doi.org/10.1016/j.ijggc.2011.05.001 Bottinga, Y., 1968. Calculation of fractionation factors for carbon and oxygen isotopic exchange in the system calcite-carbon dioxide-water. J. Phys. Chem. 72, 800–808, http://dx.doi.org/10.1021/j100849a008 Chadwick, A., Noy, D.J., Lindeberg, E., Arts, R.J., Eiken, O., Williams, G.A., 2006. Calibrating reservoir performance with time-lapse seismic monitoring and flow simulations of the Sleipner CO2 plume. In: Proceedings of the 8th International Conference on Greenhouse Gas Control Technologies, Trondheim, Norway, 19–22 June 2006. Oxford, Elsevier, pp. 1–6. Clark, I.D., Fritz, P., 1997. Environmental Isotopes in Hydrogeology. Lewis Publishers, Boca Raton, New York. Coplen, T.B., Kendall, C., Hopple, J., 1983. Comparison of stable isotope reference samples. Nature 302, 236–238, http://dx.doi.org/10.1038/302236a0 de Caritat, P., Hortle, A., Raistrick, M., Stalvies, C., Jenkins, C., 2012. Monitoring groundwater flow and chemical and isotopic composition at a demonstration site for carbon dioxide storage in a depleted natural gas reservoir. Appl. Geochem. 30, 16–32, http://dx.doi.org/10.1016/j.apgeochem.2012.05.005 Emberley, S., Hutcheon, I., Shevalier, M., Durocher, K., Mayer, B., Gunter, W.D., Perkins, E., 2005. Monitoring of fluid–rock interaction and CO2 storage through produced fluid sampling at the Weyburn CO2 -injection enhanced oil recovery site, Saskatchewan, Canada. Appl. Geochem. 20, 1131–1157, http://dx.doi.org/ 10.1016/j.apgeochem.2005.02.007 Fortier, S.M., 1994. An on-line experimental/analytical method for measuring the kinetics of oxygen isotope exchange between CO2 and saline/hypersaline salt solutions at low (25–50 ◦ C) temperatures. Chem. Geol. 116, 155–162, http://dx. doi.org/10.1016/0009-2541(94)90164-3 Halas, S., Szaran, J., Niezgoda, H., 1997. Experimental determination of carbon isotope equilibrium fractionation between dissolved carbonate and carbon dioxide. Geochim. Cosmochim. Acta 61, 2691–2695, http://dx.doi.org/10.1016/ S0016-7037(97)00107-5 Haszeldine, R.S., 2009. Carbon capture and storage: how green can black be? Science 325, 1647–1652, http://dx.doi.org/10.1126/science.1172246 Hoefs, J., 2009. Stable Isotope Geochemistry. Springer. Horita, J., Cole, D.R., Wesolowski, D.J., 1995. The activity-composition relationship of oxygen and hydrogen isotopes in aqueous salt solutions: III. Vapor–liquid water equilibration of NaCl solutions to 350 ◦ C. Geochim. Cosmochim. Acta 59, 1139–1151, http://dx.doi.org/10.1016/0016-7037(95)00031-T Johnson, G., Mayer, B., 2011. Oxygen isotope exchange between H2 O and CO2 at elevated CO2 pressures: implications for monitoring of geological CO2 storage. Appl. Geochem. 26, 1184–1191, http://dx.doi.org/10.1016/j.apgeochem.2011. 04.007 Johnson, G., Mayer, B., Shevalier, M., Nightingale, M., Hutcheon, I., 2011. Tracing the movement of CO2 injected into a mature oilfield using carbon isotope abundance ratios: the example of the Pembina Cardium CO2 monitoring project. Int. J. Greenh. Gas Control 5, 933–941, http://dx.doi.org/10.1016/j. ijggc.2011.02.003 Kharaka, Y., Cole, D.R., Hovorka, S.D., Gunter, W.D., Knauss, K.G., Freifeld, B.M., 2006. Gas–water–rock interactions in Frio formation following CO2 injection: implications for the storage of greenhouse gases in sedimentary basins. Geology 34, 577–580, http://dx.doi.org/10.1130/g22357.1

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