Water–rock interactions during a CO2 injection field-test: Implications on host rock dissolution and alteration effects

Chemical Geology 265 (2009) 227–235

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Chemical Geology j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / c h e m g e o

Water–rock interactions during a CO2 injection field-test: Implications on host rock dissolution and alteration effects N. Assayag a,b,⁎, J. Matter c, M. Ader a,b, D. Goldberg c, P. Agrinier a,b a b c

Laboratoire de Géochimie des Isotopes Stables, Institut de Physique du Globe de Paris & Université Paris 7-UMR CNRS 7154, 4, place Jussieu, 75252 Paris cedex 05, France Centre de Recherches sur le Stockage Géologique du CO2 IPGP-Total-Schlumberger, IPGP, 4, place Jussieu 75252 Paris cedex 05, France Lamont-Doherty Earth Observatory, Columbia University, 61 Route 9W, Palisades, N.Y, 10964, USA

a r t i c l e

i n f o

Article history: Received 18 February 2008 Received in revised form 2 February 2009 Accepted 12 February 2009 Editor: J. Fein Keywords: Stable isotopes (δ13C, δ18O) Carbon geological storage CO2–water–rock interaction

a b s t r a c t We investigated the nature and rates of in-situ CO2–fluid–rock reactions during an aqueous phase CO2 injection test. Two push–pull test experiments were performed at the Lamont–Doherty Earth Observatory test site (New York, USA): a non reactive control test without CO2 addition and a reactive test with CO2 equilibrated with the injected solution at a partial pressure of 1.105 Pa. The injected solution contained chemical and isotopic conservative tracers (NaCl and 18O) and was injected in an isolated and permeable interval at approximately 250 m depth. The injection interval was located at the contact zone between the Palisades sill (chilled dolerite) and the underlying metamorphic Newark Basin sediments and the injected solution incubated within this interval for roughly 3 weeks. Physico-chemical parameters were measured on the surface (pH, temperature, electrical conductivity) and water samples were collected for chemical (Dissolved Inorganic Carbon — DIC, major ions) as well as for isotopic (δ13CDIC, δ18O) analyses. For the control test, post-injection chemical and isotopic compositions of recovered water samples display mixing between the background water and the injected solution. For the reactive CO2 test, observed δ13CDIC and DIC both increase, and enrichment in Ca2+, Mg2+, K+ allow for quantification of the chemical pathways through which aqueous CO2 and subsequent H2CO3 were converted into HCO− 3 . Dissolution of carbonate minerals was the dominant H2CO3 neutralization process (≈ 52 ± 7%), followed by cation exchange and/or dissolution of silicate minerals (≈45 ± 10%, for both processes), and to a minor extent, mixing of the injected solution with the formation water (≈3 ± 1%). The results confirm the rapid dissolution kinetics of carbonate minerals compared to those of basic silicate minerals. However, our results remain marked by uncertainties due to the natural variability of the background water composition, in mass balance calculations. These experiments imply that the use of accurate DIC measurements can quantify the relative contribution of CO2– fluid–rock reactions and evaluate the geochemical trapping potential for CO2 storage in reactive reservoir environments. © 2009 Elsevier B.V. All rights reserved.

1. Introduction The recovery of CO2 emissions and their injection into deep geological reservoirs is one option envisaged to reduce their release to the atmosphere and to mitigate global warming (e.g. Bachu et al., 1994; Bachu and Adams, 2003; Chow et al., 2003; IPCC, 2005). Once the CO2 is injected into deep geological reservoirs, it undergoes a sequence of geochemical reactions with the host formation and its contained water, which are controlled by the composition of the formation water, pressure, temperature and rock mineralogy (Gunter et al., 2004; Rochelle et al., 2004). Understanding of these geochem-

⁎ Corresponding author. Laboratoire de Géochimie des Isotopes Stables, Institut de Physique du Globe de Paris & Université Paris 7 — UMR CNRS 7154, 4, place Jussieu, 75252 Paris cedex 05, France. Tel.: +33 1 44 27 28 11; fax: +33 1 44 27 28 30. E-mail address: [email protected] (N. Assayag). 0009-2541/$ – see front matter © 2009 Elsevier B.V. All rights reserved. doi:10.1016/j.chemgeo.2009.02.007

ical reactions is essential to establish the nature and rate of CO2 geochemical trapping in a storage reservoir. Until now, mass transfer and reactive transport models have been used to investigate CO2–fluid-rock reactions in response to CO2 injection in geological reservoirs (e.g. Gunter et al., 1993, 1997, 2000; Johnson et al., 2004; Xu et al., 2004; Gaus et al., 2005; Lagneau et al., 2005; White et al., 2005; Xu et al., 2005; Zwingmann et al., 2005; Andre et al., 2007; Bickle et al., 2007; Gherardi et al., 2007; Xu et al., 2007). These studies have suggested that the dissolution of host rocks would geochemically trap injected CO2. The geochemical reactivity of rocks have also been studied in many laboratory experiments at pressure and temperature conditions relevant for CO2 storage and, thermodynamic and kinetic parameters have been defined for some of the important geochemical reactions (e.g. Gunter et al.,1997; Shiraki and Dunn, 2000; Kaszuba et al., 2003; Kirste et al., 2004; Brosse et al., 2005; Kaszuba et al., 2005; Wigand et al., 2008). Such numerical and laboratory approaches suffer some

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drawbacks such as the appropriate representation of the geological and hydrological storage reservoir and, the complex coupling and hierarchy of physico-chemical processes (Marini, 2007). CO2–fluid-rock reactivity experiments performed in-situ, within natural reservoir rocks, provides a complementary approach that partly circumvents problems associated with laboratory and modelling approaches. They are critical for validating laboratory studies and useful to check the consistency between modelling predictions and experimental studies. A number of previous demonstration and pilot CO2 injection projects have been aimed at testing and developing technology for evaluating the fate of injected CO2 in the field (Emberley et al., 2004; Raistrick et al., 2006; Kharaka et al., 2006a,b; Hovorka et al., 2006; Mito et al., 2008). Such field studies have proposed active geochemical trapping of injected CO2 through dissolution of both carbonate and Ca–Mg silicate minerals. Rocks rich in Ca and Mg silicate minerals (e.g. basalts) are favorable for CO2 storage because of their high reactivity and mineral carbonation potential (McGrail et al., 2006; Matter et al., 2007; Oelkers et al., 2008). Matter et al. (2007) demonstrated the in-situ reactivity of basalt and metasedimentary rocks using small-scale push–pull CO2 injection experiments at the Lamont–Doherty test site in New York, USA. They suggest that the basalt rocks at this site are more reactive than metasedimentary rocks, which mainly consist of quartzite and metamorphic shales, and provide quantitative estimation of in-situ bulk rock dissolution rates. In this study, we report subsequent experiments at the same test site and further quantify the processes leading to CO2 neutralization after injection. These experiments offer insight into geochemical processes at low CO2 partial pressure, as it would occur far field from a point of CO2 injection at supercritical or higher pressures. Similar to previous field-scale CO2 injection experiments, we also use stable carbon and oxygen isotopes (δ13CDIC, δ18O) in combination with chemistry data to describe in-situ CO2–fluid-rock reactions and to be sensitive tracers of the CO2 reactivity (Emberley et al., 2004; Kharaka et al., 2006a,b; Raistrick et al., 2006). 2. Methodology

Table 1 Characteristics of the small push–pull experiments for the control test and the CO2 test.

Conservative tracers Volume of 18O spike (ml) Mass of NaCl (g) Reactive component PCO2 in equilibrium with the injected solution (105 Pa) Push phase Test solution (l) Push phase duration (h) Injection rate (l min− 1) Incubation period (days) Pull phase Background water/test solution (l) Pull phase duration (h) Pumping rate (l min− 1)

Control test

CO2 test

May 2005

July 2005

30 210 –

30 210 1

1299 3 7 28

1320 3 7 24

9300 16.3 9.6

20,700 34 10

over the same hydraulically isolated interval and the same time period. Detailed characteristics of both tests are summarized in Table 1. Both the control and reactive tests each lasted for an incubation period of 3 weeks, during which time the injected fluid mixed with the nature groundwater. After the incubation period, the injected solution/background water mixture was pumped back using a submersible pump and passed through a measurement cell in which electrical conductivity, temperature and pH were continuously measured. The elapsed pumping time and the extracted fluid volume were recorded automatically, and samples were collected at incremental periods from the discharge stream (from one sample per half-hour to one sample twice a day). Note that prior to the injection, samples of the background water within the isolated interval, and of the injected solution, were collected for chemical and isotopic analysis. CO2 partial pressures of the extracted water samples were calculated using PhreeqC program (Parkhurst and Appelo, 1999) and ranged from 10 Pa to 1.7.104 Pa; they were below 1.105 Pa (total atmospheric pressure) showing no potential CO2 outgassing during the pumping back.

2.1. Single well push–pull test 2.2. Analytical methods Single well push–pull tests were conducted in the contact zone between the Palisades sill and the underlying Newark Basin sediments at the Lamont–Doherty Earth Observatory test site in Palisades, New York, USA. Petrophysical and hydrological characterizations of the target injection interval as well as the overlying and underlying formations have been reported previously (Burgdorff and Goldberg, 2001; Goldberg and Burgdorff, 2005; Matter et al., 2006). At this site, the Palisades sill is 230 m thick and consists of dolerite rich in plagioclase and pyroxene. The contact zone between the dolerite and the underlying sediments is approximately 10 m thick and is characterized by chilled dolerite and contact-metamorphosed sedimentary rocks and by an increased bulk permeability. Below the contact zone, alternating layers of siltstones and mudstones of the Newark Basin group sediments are dominant (Fig. 1). Two single well push–pull tests have been performed previously at this site, and are described by Matter et al. (2007). For this study, two injections were conducted following the same methodology. In the first test, a control experiment, we introduced about 1300 l of an aqueous solution into a hydraulically isolated interval at 232–240 m depth (the contact zone) over a time period of 3 h. The injected solution was prepared in a polyethylene tank at the wellhead and consisted of formation water pumped from the contact zone (called background water, BW), spiked with oxygen-18 isotopes (30 ml, H18 2 O enrichment: 98.2%) and sodium chloride (210 g) both used as conservative tracers. In the second experiment, the reactive CO2 test, conducted 30 days later, we introduced about 1300 l of the same injected solution equilibrated with 1.105 Pa CO2 partial pressure (PCO2)

A total of 12 water samples were collected in the control test and 32 water samples in the reactive CO2 test. Sterilized polyethylene vials were used in sampling for ion analyses and Labco Exetainer glass tubes were used in sampling for isotopic analysis. All fluid samples were filtered with 0.2 mm nylon filters. Water samples used for the major and minor ions measurements were stored in 1 M HNO3 preacidified vials. Major and minor cations (Ca2+, Mg2+, Na+, K+, åSi) were measured by Inductively Coupled Plasma-Atomic Emission Spectrometry ICP-AES (Optima 3000) at the XRAL laboratory in Toronto, Canada. Multiple (n = 4) sample analysis resulted in an average reproducibility of ±0.04 mmol l− 1. Anions (Cl− and HCO− 3 ) were analyzed by capillary electrophoresis using a Water Capillary Ion Analyzer with an analytical precision of about 5%, at the laboratory of water geochemistry (IPGP, France) (Weinberger, 2000). Dissolved Inorganic Carbon (DIC) concentration was calculated from HCO− 3 measurements by electrophoresis and pH data, for water samples with pH greater than 6.3 (Weinberger, 2000). For water samples with pH lower than 6.3, like the injected solution with its low pH (≈4.8), DIC concentration was measured using Gas Chromatography-Isotope Ratio Mass Spectrometer AP2003 (GC-IRMS) (Assayag et al., 2006) since its HCO− 3 concentration was too low to be detected by electrophoresis. Measurements of δ13C on both fluid and rock samples were made using an aliquot of fluid or powdered rock material for reaction with H3PO4; the released CO2 was analyzed using the GC-IRMS. The measured carbon isotope compositions were calibrated using a set

(a) Calculated with Eq. (1). (b) Calculated from the average value of repeated analysis of background waters sampled before the control and before CO2 test. Average ((1) + (2) + (3) + (4)) (Tables 2 and 4).

− 6.0 − 7.1 − 8.3 − 8.8 − 8.9 − 8.9 − 8.9 − 9.0 − 9.0 − 8.8 1.2 Control_06_13_05_1 Control_06_13_05_2 Control_06_13_05_3 Control_06_14_05_1 Control_06_14_05_2 Control_06_15_05_1 Control_06_15_05_2 Background water (BW)(1) Background water (BW)(2) Background water (BW)(b) Injected solution (IS)

144.3 577.2 2164.5 4324.5 6198.5 6909.3 9339.4 – – – –

9.2 9.3 9.3 9.4 9.3 9.4 9.4 9.3 9.3 9.3 9.4

527 474 444 429 423 423 423 493 – 439 877

17.3 17.3 16.7 17.3 16.5 16.0 15.1 14.9 14.7 15.3 16.4

1.53 1.50 1.46 1.46 1.67 1.56 1.63 1.37 1.36 1.23 1.49

0.13 0.11 0.10 0.10 0.09 0.09 0.09 0.15 0.14 0.16 0.13

0.03 0.02 0.03 0.03 0.03 0.03 0.03 0.03 0.03 0.03 0.02

0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01

4.31 4.12 3.70 3.72 3.71 3.63 3.66 3.47 3.47 3.29 6.11

0.27 0.26 0.26 0.26 0.26 0.26 0.25 0.31 0.31 0.33 0.31

2.52 2.14 1.66 1.54 1.32 1.33 1.29 1.29 1.37 1.47 4.17

− 18.3 − 18.2 − 17.4 − 17.3 − 17.0 − 17.0 − 17.1 − 17.4 − 17.4 − 17.8 − 17.6

X(a) 18O ‰

δ18O δ13CDIC

‰ mmol l− 1

Cl− ∑Si

mmol l− 1 mmol l− 1

Na+ Mg2+

mmol l− 1 mmol l− 1

K+ Ca2+

mmol l− 1 µS.cm− 1

DIC Temperature

°C

Conductivity pH

Pumped volume (l) Samples (EWS)

The chemical and isotopic results from the CO 2 test are summarized in Table 4. Similar to the control test, NaCl and 18O tracers were added to the injected solution yielding a Cl− ion

Table 2 Control test: physico-chemical parameters, chemical, isotopic data and mixing fractions (X18O, XCl−).

3.2. CO2 test

Control_ sampling date_number

3.1. Control test

mmol l− 1

3. Results

Chemical and isotopic results from the control test are summarized in Table 2. The addition of NaCl and 18O tracers resulted in large differences in Cl− ion and δ18O levels between the background water and the injected solution. The Cl− ion concentration was of 4.17 mmol l− 1 and δ18O value was of 1.2‰ in the injected solution compared to 1.47 mmol l− 1 and −8.8‰, respectively, in the background water. During the pull phase of the experiment, Cl− ion concentrations and δ18O values decreased from 2.52 to 1.29 mmol l− 1 and from −6.0 to −9.0‰, respectively. These variations can be fully explained by mixing between the background water and the injected solution. The concentration of major and minor ions (Ca2+, Mg2+, K+, 13 ∑Si = H4SiO4 + H3SIO− 4 ), DIC and δ CDIC values from the extracted water samples remained close to those of the injected solution and the background water. The DIC concentrations of the extracted water samples ranged from 1.46 to 1.67 mmol l− 1 and the corresponding δ13CDIC values ranged between −18.3‰ and −17.0‰. DIC concentrations of the background water and injected solution were 1.23 mmol l− 1 and 1.49 mmol l− 1, respectively. Values of δ13CDIC of the background water and injected solution were −17.8‰ and −17.6‰, respectively. The consistency of these results indicates a non reactive water–rock system. Slight compositional differences in the extracted water samples relative to those of the injected solution and background water are likely due to uncertainties introduced by single well push–pull tests, the natural variability of the background water and the analytical uncertainties. The magnitudes of these uncertainties are estimated by two different methods and indicate the significance levels above which compositional changes may be attributed to chemical reactivity. These uncertainties are presented in Table 3. In the first method, significance levels are calculated by the difference between maximum and minimum values for the extracted water samples measured during the control test. In the second method, significance levels are calculated by the difference between the mean value of an extracted water sample and the background water. Both uncertainties determined by both methods agree fairly well.

0.28 ± 0.04 0.17 ± 0.04 0 0.05 ± 0.04 0.00 ± 0.04 – – – – – – –

(a) XCl−

of CaCO3 standards to provide δ13C versus the PDB (Pee Dee Belemnite) standard scale (Mook, 2000), and yielding δ13C = [(13C/ 12 C)sample / (13C/12C)ref − 1] ⁎ 1000‰. The analytical accuracy (1σ) was better than ± 0.1‰ for the δ13CDIC of fluid samples and better than ± 0.3‰ for the δ13CCaCO3 of rock samples. Further details about the method are discussed by Assayag et al. (2006). The oxygen isotope compositions of fluid samples were also measured using the GC-IRMS (AP2003) after equilibration of an aliquot of water sample with a He+CO2 gas mixture. The measurements were calibrated using a set of water standards to provide δ18O versus the SMOW (Standard Mean Ocean Water) standard scale. Measurements of δ18O were determined with an analytical accuracy (1σ) of better than ±0.1‰. Because sources and sinks of oxygen during mineral dissolution or other in-situ processes add or consume less than a few tens of mmol per liter of oxygen, and the injected solution contains roughly 55,500 mmol l− 1 or four orders of magnitude more oxygen, δ18O isotopes can be effectively used as a conservative tracer. Even though oxygen content may be enriched by as much as 42.2‰ relative to the injected solution, the maximum δ18O shift will be at most 0.1‰, within the precision of δ18O measurements.

229 0.39 ± 0.1 0.25 ± 0.1 0.07 ± 0.1 0.03 ± 0.1 – – – – – – –

N. Assayag et al. / Chemical Geology 265 (2009) 227–235

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N. Assayag et al. / Chemical Geology 265 (2009) 227–235

Table 3 Significance levels determined from the control test data. Ca2+

DIC mmol l Method I Method II

0.210 0.314

−1

mmol l

K+ −1

0.039 0.058

mmol l

Mg2+ −1

0.008 0.001

mmol l

δ13CDIC

∑Si −1

0.006 0.004

mmol l

−1

0.0200 0.0001

‰ 0.400 0.329

Method I: The significance level for a component C is the difference between the maximum and the minimum values of the component C measured in the extracted water samples. Method II: The significance level for a component C is the absolute value of the difference between the mean value of the component C measured in the extracted water sample and the value of the component C measured in the background water. Due to the spiking with NaCl and 18O spike, no value can be calculated for Na+, Cl−, δ18O data.

concentration of 5.16 mmol l− 1 and a δ18O value of 2.1‰ in the injected solution whereas in the background water, Cl− ion concentration was 1.47 mmol l− 1 and the δ18O value was −8.8‰. During the pull phase of this experiment, Cl− ion concentrations and δ18O values decreased from 2.64 to 1.37 mmol l− 1 and from −5.2 to −9.0‰, respectively. These variations were due to mixing processes between the background water and the injected solution within the aquifer. Based on the conservative tracers (i.e., Cl− and 18O), the total recovery ratio of the injected to recovered fluid volume was estimated at 25 ± 5%. This ratio is computed using the amount of chloride pumped during the pull phase of the experiment divided by the amount of chloride contained in the injected solution. Unlike the control test, however, DIC concentrations and δ13CDIC values changed as a result of equilibration of the injected solution. At a CO2 partial pressure of 1.105 Pa, DIC concentrations increased from 1.2 mmol l− 1 in the background water to 41.4 mmol l− 1 in the injected solution, and δ13CDIC values decreased from −17.8 to −51.1‰. Low δ13C levels in the injected solution are not fully understood, but may be explained by oxidation of the CO2 gas loaded originally into the injected solution. This gas was commercially obtained and may contain 13C depleted methane or favor 12C depleted CO2 by industrial membrane diffusion processes.

During the pull phase of the reactive experiment, the chemical and isotopic compositions of the extracted water samples (Ca2+, Mg2+, Na+, K+, ∑Si, DIC and δ13CDIC) evolved during the first three days of pumping until they nearly reached the pre-injection background water levels the fourth day. Measured DIC concentrations decreased from 18.91 to 1.61 mmol l− 1, and δ13CDIC values increased from −41.5 to −22.4‰. Calcium and sodium ion concentrations displayed sharp decreases from 2.94 to 0.14 mmol l− 1 and from 6.14 to 3.37 mmol l− 1, respectively. Saturation indices for calcite ranged between −0.99 and −0.09, and indicated that the first fifteen extracted water samples were all undersaturated with respect to calcite (In Table 5, saturation indexes were only reported for the first six extracted water samples). The 4 days (with 8 h of pumping per day) of the pull phase were insufficient to reach accurately the pre-injection levels of DIC and δ13CDIC. Additional days of pumping back would have been necessary. In addition to mixing, these results from the chemical and isotopic analysis of the recovered water samples indicate that other geochemical reactions and processes occurred within the aquifer following injection of CO2-saturated solution. These reactions are discussed below. 3.3. Mixing end-members The chemical and isotopic compositions of the background water display variations for Na+ (14%), Cl− (35%), DIC (29%), δ13CDIC (0.9‰) and δ18O (0.5‰), over the 4-months-period during which the two push–pull tests were conducted. These fluctuations are natural and inherent to shallow aquifers stressed by pumping. They may also reflect the seasonal and temporal trends of the recharge waters, the hydrodynamic of the aquifer and the biological activity (mainly for DIC and δ13CDIC). Indeed, there is evidence that the groundwater within the injection interval is relatively young (unpublished data), and hence, this variability illustrates the difficulty to define the background water end-member. Due to such inherent natural fluctuations, for both tests, we have chosen to use mean values for the chemical and isotopic compositions of the background water. We average values of repeated analyses of background water — two

Table 4 CO2 test: physico-chemical parameters, chemical, isotopic data. Samples (EWS) CO2_ sampling date_number CO2_07_25_05_1 CO2_07_25_05_2 CO2_07_25_05_3 CO2_07_25_05_4 CO2_07_25_05_5 CO2_07_25_05_6 CO2_07_25_05_7 CO2_07_25_05_8 CO2_07_25_05_9 CO2_07_25_05_10 CO2_07_25_05_11 CO2_07_26_05_1 CO2_07_26_05_2 CO2_07_26_05_3 CO2_07_26_05_4 CO2_07_26_05_5 CO2_07_27_05_1 CO2_07_27_05_2 CO2_07_27_05_3 CO2_07_28_05_1 CO2_07_28_05_2 Background water(BW)(3) Background water(BW) (4) Background water(a) (BW) Injected solution (IS)

Pumped volume (l)

pH

210.2 330.3 680.7 1131.1 1381.4 1931.9 2382.4 2983.0 3533.5 4134.1 4634.6 5432.9 6172.4 7207.7 7799.3 8785.3 11,039.3 12,402.8 13,311.8 15,875.9 17,227.2 – – – –

6.5 6.5 6.6 6.7 6.8 6.9 7.0 7.1 7.2 7.4 7.4 7.4 7.8 8.0 8.2 8.4 8.6 8.8 8.9 9.0 9.0 9.4 9.2 9.3 4.8

Conductivity µS∙cm 1143 1195 1010 852 783 685 636 579 559 527 512 478 475 466 458 447 440 443 419 435 429 411 412 439 710

−1

Temperature

Ca2+

DIC

°C

mmol l

15.5 15.7 15.9 15.7 15.8 16.8 16.6 16.5 16.1 16.0 16.0 16.8 17.2 16.8 16.1 15.8 17.2 16.7 16.3 15.8 15.8 16.4 – 15.3 20.5

18.38 18.91 13.26 10.29 8.33 5.21 2.32 3.42 3.02 2.54 2.30 2.38 1.89 1.58 1.57 1.93 1.77 1.71 1.68 1.65 1.61 1.06 1.13 1.23 41.4

−1

mmol l 2.94 2.70 1.88 1.47 1.22 0.91 0.75 0.63 0.54 0.46 0.42 0.45 0.35 0.28 0.26 0.23 0.19 0.17 0.16 0.15 0.14 0.20 0.14 0.16 0.20

K+ −1

mmol l 0.05 0.05 0.04 0.04 0.04 0.04 0.03 0.03 0.03 0.02 0.03 0.03 0.03 0.02 0.02 0.02 0.02 0.02 0.02 0.02 0.02 0.03 0.02 0.03 0.03

Mg2+ −1

mmol l 0.11 0.11 0.09 0.08 0.07 0.05 0.04 0.03 0.03 0.03 0.02 0.02 0.02 0.02 0.02 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.02

Na+ −1

Cl−

∑Si −1

−1

mmol l

mmol l

mmol l

5.83 6.14 5.74 5.32 5.02 4.57 4.28 4.05 3.92 3.80 3.75 3.82 3.62 3.57 3.76 3.56 3.4 3.38 3.38 3.4 3.37 3.18 3.05 3.29 5.76

0.39 0.37 0.34 0.31 0.30 0.27 0.26 0.25 0.25 0.25 0.25 0.25 0.24 0.23 0.24 0.23 0.23 0.22 0.23 0.23 0.22 0.40 0.31 0.33 0.31

2.64 2.55 2.16 1.97 1.88 1.80 1.71 1.63 1.59 1.54 1.49 1.47 1.42 1.38 1.38 1.43 1.41 1.41 1.39 1.39 1.37 1.74 – 1.47 5.16

−1

δ13CDIC

δ18O

‰

‰

− 41.5 − 41.4 − 40.4 − 39.5 − 38.4 − 36.4 − 35.2 − 33.7 − 32.5 − 31.7 − 31.3 − 31.2 − 29.7 − 28.0 − 27.0 − 26.2 − 25.0 − 23.9 − 23.5 − 23.3 − 22.4 − 18.3 − 18.1 − 17.8 − 51.1

− 5.2 − 5.3 − 6.4 − 7.0 − 7.5 − 7.9 − 8.3 − 8.3 − 8.6 − 8.6 − 8.8 − 8.7 − 8.7 − 8.8 − 8.8 − 8.9 − 8.9 − 9.1 − 9.0 − 9.0 − 9.0 − 8.5 − 8.6 − 8.8 2.1

(a) Calculated from the average value of repeated analysis of background waters sampled before the control and before the CO2 test. Average ((1) + (2) + (3) + (4)) (Tables 2 and 4).

N. Assayag et al. / Chemical Geology 265 (2009) 227–235

231

Table 5 CO2 test: mixing fractions (X18O, XCl−), chemical mass balances (Δ = CEWS − Cmix), calcite saturation index ISCaCO3. Samples (EWS) CO2_date_number CO2_07_25_05_1 CO2_07_25_05_2 CO2_07_25_05_3 CO2_07_25_05_4 CO2_07_25_05_5 CO2_07_25_05_6 a b c d e

X18Oa

XCl−a

ΔCa2+b

ΔK+b

mmol l

mmol l

0.33 ± 0.04 0.32 ± 0.07 2.77 0.32 ± 0.04 0.29 ± 0.07 2.53 0.22 ± 0.04 0.19 ± 0.07 1.72 0.17 ± 0.04 0.14 ± 0.07 1.31 0.12 ± 0.04 0.11 ± 0.07 1.06 0.08 ± 0.04 0.09 ± 0.07 0.75

−

1

−

1

2.00E− 02 2.30E− 02 1.50E− 02 1.20E− 02 1.20E− 02 6.70E− 03

ΔMg2+b

ΔNa+b

mmol l

mmol l

−

1

9.40E− 02 9.70E− 02 7.70E− 02 6.00E− 02 5.10E− 02 3.60E− 02

−

1.72 2.04 1.89 1.62 1.43 1.08

1

Δ(∑Si)b

DICmixc

mmol l

mmol l

0.04 0.03 0.00 − 0.03 − 0.04 − 0.06

−

1

ΔDICb −

1

14.49 14.09 10.47 7.66 6.05 4.44

mmol l

δ13Cmixc Δδ13CDICb δ13CDIC−adde ISCaCO3

ΔDIC⁎d −

1

3.88 ± 1.0 4.78 ± 1.0 3.19 ± 0.7 2.43 ± 0.8 2.31 ± 0.8 0.90 ± 1.1

mmol l

−

3.4 ± 0.5 3.4 ± 0.5 2.4 ± 0.5 1.9 ± 0.5 1.4 ± 0.5 ≈1

1

‰

‰

‰

49.1 − 49.1 − 47.8 − 46.5 − 44.8 − 41.6

7.7 ± 0.3 7.7 ± 0.3 7.5 ± 0.5 7.2 ± 0.7 6.7 ± 1.3 5.2 ± 3.4

12.7 18.6 − 16.6 − 16.0 − 20.9 − 12.8

−

0.28 0.30 − 0.44 − 0.50 − 0.53 − 0.72

−

−

−

−

Calculated with Eq. (1) and data from Table 4. Calculated with Eq. (6) or Eq. (5) and data from Table 4. Calculated with Eq. (2) or Eq. (3) and data from Table 4. Conditional ΔDIC, see Section 4.2. Calculated with Eq. (7).

sample sets prior to the control test (Table 2 (a, b)) and two sample sets prior to the CO2 test (Table 4 (c, d)). The chemical and isotopic compositions of the injected solution end-members are reported in Tables 2 and 4.

3.4. Chemical and carbon isotopic compositions of crushed rock samples Fig. 1 presents the total inorganic calcium carbonate content and its carbon isotope composition (δ13CCaCO3), measured on crushed drill cuttings, collected as the hole was being drilled. CaCO3 content varies from 0.05 to 3.57 wt.%, where the δ13CCaCO3 values range between −14‰ and − 6‰.

4. Discussion 4.1. CO2 reactivity The chemical and isotopic compositions of the extracted water samples (EWS) highlight the reactive behaviour of the CO2 test (i.e. mixing + CO2–water–rock interactions) and the non-reactive behaviour of the control test (i.e. only mixing). In order to assess the CO2 reactivity and the H2CO3 neutralization capacity during the CO2 test, we first calculated the mixing proportions between the injected solution (IS) and the background water (BW) using mass balance equations for the conservative tracers: ½TracerEWS = X ½TracerIS + ð1 − XÞ½TracerBW

ð1Þ

where [Tracer] is the measured Cl− ion concentration or δ18O values and X is the fraction of injected solution in the extracted water sample. For the first six extracted water samples of the CO2 test, X ranges from 0.33 to 0.08 for δ18O data and from 0.32 to 0.09 for Cl− ion data (Table 5). The remaining samples of the CO2 test will not be considered further since their fractions X are lower than their uncertainties. The good agreement of X calculated using either one of the conservative tracers is highlighted in Fig. 2, for the CO2 test. For the following analysis, we chose to use the X defined with δ18O data (X18O), since δ18O data was measured with a higher accuracy than Cl− ion concentration. Once X values have been established for the first six extracted water samples, their theoretical δ13CDIC and DIC values

Fig. 1. Total inorganic calcium carbonate content (wt.%) and δ13CCaCO3 values (‰): profile along the borehole above and below the target injection interval. Formation name and rock nature is indicated on the left part of the diagram (in bold and italic, respectively). CaCO3 weight % is given close to each data points. The dotted line delimits the interval of depth in which the push–pull tests were performed. (The abbreviation “n–m” means that there was no measurement for those rock samples since their inorganic calcium carbonate contents were too low to be detected).

− Fig. 2. Comparison of X18O and XCl , mixing proportions between background water and injected solution estimated respectively with δ18O and Cl− concentration data (full square: CO2 test; open square: control test). For the CO2 test, the distance between the best fit line (solid line) and the 1:1 line (dashed line) is smaller than the error bars thus confirming the good consistency between these two methods of evaluation of mixing − proportions. For the control test, the agreement between X18O and XCl is moderate due to the spatial and temporal variabilities of the background water composition.

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with measured δ13CCaCO3 values of − 9 to − 7‰ within the injection interval. Note that using DIC system in this manner does not require prior knowledge of the fractionation of carbon isotopes between the 2− species of the DIC (H2CO3, HCO− and CaCO3). 3 , CO3

4.2. Uncertainty analysis The procedure used to compute the extent of the CO2 reactivity is based on a chain of equations (Eqs. (1) to (5)) that propagates and accumulates the uncertainties of the different parameters (δ18OIS, δ18OBW, δ18OEWS, δ13CDIC–IS, δ13CDIC–BW, δ13CDIC–EWS, DICIS, DICBW, DICEWS) onto the final estimations of ΔDIC and δ13CDIC-add. To investigate further, we conducted a Monte-Carlo analysis which

Fig. 3. Relationship between δ13CDIC and DIC concentrations. The grey thick curve is the mixing hyperbola between the two end-members. Large black circles, full and open, are for the injected solution and the background water, respectively. For the first six extracted water samples of the CO2 test, black crosses refer to the theoretical mixing data (δ13CDIC–mix and DICmix) calculated using mass balance Eqs. (2) and (3). Black dots refer to the measured data (Table 4). Arrows link theoretical mixing data and measured data for the first six extracted samples of the CO2 test.

(δ13CDIC–mix and DICmix) due to pure mixing can be calculated using the following mass balance equations: DICmix = X180 · DICIS + ð1 − X180 Þ · DICBW 13

13

ð2Þ 13

δ CDIC–mix DICmix = X180 · DICIS δ CDIC–IS + ð1 − X180 Þ · DICBW δ CDIC–BW

ð3Þ which yield: " 13

δ CDIC–mix =

 3 2 # DICBW DICIS δ13 CDIC–BW − δ13 CDIC − IS DICIS δ13 CDIC–IS − DICBW δ13 CDIC–BW 5: +4 DICIS − DICBW DICmix ðDICIS − DICBW Þ

ð4Þ 13

In a δ CDIC versus DIC diagram, the theoretical mixing data values plot along a mixing hyperbola. However, the measured data points (δ13CDIC–EWS and DICEWS) are located above the theoretical mixing hyperbola, indicating the occurrence of additional chemical processes besides pure mixing (Fig. 3). Differences in δ13CDIC values and DIC concentrations between the values measured in the extracted water samples and the theoretical mixing values yield the excesses in 13C (Δ13C) and DIC (ΔDIC): 13

13

13

Δδ CDIC = δ CDIC–EWS − δ CDIC–mix

ð5Þ

ΔDIC = DICEWS − DICmix :

ð6Þ

Eqs. (5) and (6) yield positive values for the first six extracted water samples (Table 5), indicating that the reactivity of the injected solution increased the DIC concentration and enriched it in 13C. This is consistent with carbonate dissolution by a neutralization reaction of H2CO3. The amount of added DIC (DIC-add) is given by ΔDIC and its δ13CDIC-add is computed using the following mass balance equation:   13 13 13 δ CDICadd = DICEWS δ CDIC–EWS − DICmix δ CDIC–mix = ΔDIC:

ð7Þ

For the first six extracted water samples, the δ13CDIC-add values range between −21 and −13‰, with uncertainties of about ±10‰ (Table 5). Since these water samples were undersaturated with respect to calcite, calcite dissolution is irreversible and thus transfers DIC to all its carbon isotopes without fractionation. The range of δ13CDIC-add values can then be compared directly to the δ13C measured on the disseminated calcite. Fig. 1 illustrates reasonable agreement

Fig. 4. A. Error hyperbola for the δ13Cadd and ΔDIC of the first extracted water sample CO2_07_25_05_1 (see Table 5). It was estimated via a Monte-Carlo simulation of 2000 trials (grey dots) during which the parameters (δ18OIS, δ18OBW, δ18OEWS, δ13CDIC–IS, δ13CDIC–BW, δ13CDIC–EWS, DICIS, DICBW, DICEWS) were sampled in their respective Gaussian probability distributions. These 2000 trials define a quasi-perfect hyperbola (thick black line: δ13CDIC-add = 0.128/ΔDIC − 46.13; r2 = 0.940). The most likely δ13CDIC-add and ΔDIC values (black circle) are from Table 5. The 1σ error bars are estimated from the dispersion of the 2000 trials. The 66% confidence contour (thin line) and the 95% confidence contour (dashed line) are given. On the δ13CDIC-add axis, the range of δ13CCaCO3 values (between −6 and −10‰, see Fig. 1) measured around the target injection interval allows the range for the conditional ΔDIC values (named ΔDIC⁎ hereafter) to be defined (grey domain) using the error hyperbola. B. Error hyperbola for the δ13CDIC-add and ΔDIC of the first six extracted samples of the CO2 test (from Table 5). The 66% confidence domains are reported in dark grey together with the δ13CCaCO3 domain (light grey).

N. Assayag et al. / Chemical Geology 265 (2009) 227–235

takes into account the uncertainties of these different parameters and the error correlations introduced by Eqs. (1) to (5). For the first extracted water sample CO2_07_25_05_1, ΔDIC (3.88 mmol l− 1) appears quite well determined (σ = ±1 mmol l− 1, i.e. ±26%) whereas the corresponding δ13CDIC-add = 12.7‰ is affected by an error of σ = ±10‰. The uncertainty in DICEWS concentrations is the major source of error for more than 60% of the samples. The next most significant sources of error being the uncertainties on the DIC concentrations of the two end-members (DICIS, DICBW) of roughly 15% each. All remaining uncertainties are minor and contribute to less than 10% of the calculated error for ΔDIC and δ13CDIC-add. Developing a methodology that produces more precise DIC measurements would considerably improve the precision of ΔDIC and δ13CDIC-add used for characterizing and quantifying CO2 reactivity in this and similar experiments. In Fig. 4A, we note that the uncertainties in ΔDIC and δ13CDIC-add are highly anti-correlated, showing that ΔDIC and δ13CDIC-add are not independently determined by this set of parameters. Additional information from another input parameter would result in more precise determinations of ΔDIC or δ13CDIC-add. For example, assuming that excess of DIC comes from the dissolution of measured carbonates and incorporating δ13CCaCO3 values ranging between − 10 and −6‰, we can deduce a conditional ΔDIC value, denoted ΔDIC⁎, between 3 and 4 mmol l− 1 (i.e. σ ≈ 0.5 mmol l− 1; Fig. 4A, Table 5). Similarly, ΔDIC⁎ values are estimated for the five extracted water samples of the CO2 test. ΔDIC⁎ values are about 25% lower than ΔDIC values, and both results are discussed below (Fig. 4B, Table 5). 4.3. Major ion chemical data Using Eq. (1), we compute mass balance mixing relationships between the background water and the injected solution for Ca2+, Mg2+, K+, Na+ ions and ∑Si. Then, using Eq. (6), we compute ΔCa2+, ΔMg2+, ΔK+, ΔNa+ and Δ(∑Si) excesses mass values. For the first six extracted water samples from the CO2 test, ΔCa2+, ΔMg2+, ΔK+ and ΔNa+ are greater than their significance levels whereas Δ(∑Si) remains very close to its significance level but decreases from positive to negative (Table 5). Positive ΔCa2+, ΔMg2+, ΔK+ and ΔNa+ values confirm the dominance of rock dissolution reactions over precipitation reactions. The changes in Δ(∑Si) values, assuming they are significant, would rather suggest silicate dissolution for the first two samples and silicate precipitation for the latter three samples. The total amounts (∑) of Ca2+, Mg2+, K+, Na+ ions and DIC added by CO2 reactivity during this test were estimated from the integration of the Δ excesses mass values over the pumped volumes and corrected for a recovery rate of about 25%. They are ordered as follows (Table 6): ∑DIC ≈ 1.4 ∑Na+ ≈ 1.5 ∑Ca2+ ≈ 35 ∑Mg2+ ≈ 173 ∑K+, and indicate a higher release of DIC, Na+ and Ca2+ ions compared to Mg2+ and K+ ions. If, as suggested earlier, ΔDIC and δ13CDIC-add suggest carbonate dissolution, we expect a ΔCa2+/ΔDIC ratio of 1 for CaCO3 dissolution

233

in a closed system (Deines et al., 1974; Stumm and Morgan, 1996) and a ratio b1 if the carbonate contains magnesium. Dissolution of carbonates in an open system is not relevant for this experiment because the reservoir was confined and a finite amount of CO2 was available. For the first six extracted water samples, the weighted average of (ΔCa2+ + ΔMg2+)/ΔDIC ratio is 0.69 ± 0.26 using measured DIC concentrations, and 0.94 ± 0.32 using ΔDIC⁎ values. Although lower than expected theoretically, these ratios are roughly equal within their uncertainties. If cation exchange of Ca2+ and Mg2+ ions with Na+ ion on clay minerals is taken into account (Fletcher et al., 1984; Kleven and Alstad, 1996), the comparison of derived ratios for ΔDIC and ΔDIC⁎ values improves. Cation exchanges processes cannot be ruled out with the existing data set and would consume Ca2 + and Mg2+ ions and add Na+ ions to the solution. If ΔNa+ was added to the solution entirely by cation exchange, the (ΔCa2+ + ΔMg2+)/ ΔDIC⁎ ratio would be 1.15 ± 0.32. Hence, within the sensitivity limits of our experimental tests (PCO2 ≈ 1.105 Pa and pH ≈ 4.8), the evolution of the major ions may be explained by the dissolution of disseminated calcite and complementary cation exchange. Because of the negligible measured values of Δ(∑Si), it is difficult to determine whether dissolution of silicate minerals also occurred. During the push–pull test conducted previously at this site with higher PCO2 levels (8.105 Pa, pH ≈ 3.4; Matter et al., 2007), the computed value of ∑Ca/∑Na was 9 and the measured Δ(∑Si) was much higher than in this study, indicating the dominance of carbonate and Ca–Mg silicate mineral dissolution over cation exchange. Results from both this study and Matter et al. (2007) suggest that the extent of the dissolution of silicate minerals is much smaller than that of carbonate minerals and the variation of Ca2+ ion concentration is much greater than that of Mg2+, K+, Na+. This can be explained by the high reactivity and the rapid dissolution kinetics of carbonate minerals compared to those of Ca–Mg-rich silicate minerals (Lasaga, 1995; Sverdrup and Warfvinge, 1995; Pokrovsky et al., 2005). Moreover, silicate mineral reactivity in both push–pull tests is consistent with flow-through and batch type laboratory experiments, which show that silicate mineral reactivity increases as the pH decreases (Gunter et al., 1997; Carroll and Knauss, 2005; Golubev et al., 2005; Hänchen et al., 2006).

4.4. Mass balance of H2CO3 consumption For the CO2 injection test in this study, the amount of reacted H2CO3, 28 ± 5 mol, is given by the difference between the amount of injected H2CO3, 53 mol (injection of 1320 l ⁎ 40.2 mmol l− 1) and the amount of the re-pumped and unreacted H2CO3, 25 ± 5 mol. Thus, integrating H2CO3 concentrations over the pumped volume and correcting for the recovery rate of 25%, the total consumption of H2CO3 is computed and shown in Table 7. The consumption of H2CO3 is the consequence of several in-situ processes:

Table 6 CO2 test: Volume integrated amounts of unreacted H2CO3 and of Ca2+, K+, Mg2+, Na+, DIC added by CO2 reactivity. Samples (EWS) CO2_date_number CO2_07_25_05_1 CO2_07_25_05_2 CO2_07_25_05_3 CO2_07_25_05_4 CO2_07_25_05_5 CO2_07_25_05_6 ∑b

Pumped volume (l)

∑Ca2+

∑K+

∑Mg2+

∑Na+

∑DIC

∑DIC⁎

H2CO3a

mmol

mmol

mmol

mmol

mmol

mmol

mmol

210.2 330.3 680.7 1131.1 1381.4 1931.9

582 304 603 590 265 413 11,028

4 3 5 5 3 4 96

20 12 27 27 13 20 471

362 245 662 730 358 595 11,804

816 574 1118 1094 578 0 16,720

715 412 848 847 358 55 12,937

1494 878 1551 1319 501 576 25,280

NB: ∑x = ΔY ⁎ (Vfinal − Vinitial), where Y is either Ca, K, Na, Mg or DIC concentrations. a Calculated from pH, DIC and acidity constants (K1 = 10− 6.3 and K2 = 10− 10.3). b Corrected for 25% recovery rate.

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Table 7 H2CO3 mass balances (in mol). Injected H2CO3 Min Max Mean Sigma

53

Pumped H2CO3

25 5

Reacted H2CO3

28 5

H2CO3 consumed by mixinga

H2CO3 consumed by dissolution of carbonatesb

H2CO3 consumed by cation exchangec

0.7 0.25

12.9 16.7 14.8 2

0 11.8 5.9 3

TDScatd− 2.QDIC

Balancee

5 4.5

1.7 6

a Evaluated by the difference between the “theoretical” H2CO3 content assuming a pure mixing (Eq. (2)) and the “real” H2CO3 content calculated assuming the conservation of alkalinity and that of the DIC (no degassing, no dissolution, no precipitation). b Evaluated from QDIC and Q⁎ DIC. c Evaluated from QNa. d TDScat = [Na+] + [K+] + 2.[Ca2+] + 2.[Mg2+]. e Balance = reacted H2CO3 − (H2CO3 consumed by mixing) − (H2CO3 consumed by dissolution of carbonates) − (H2CO3 consumed by cation exchange) − (TDScat − 2.QDIC).

i) Mixing between background water and the injected solution according to H2CO3 + CO2− → 2 HCO− 3 3 . The amount of neutralized H2CO3 is evaluated by the difference between the theoretical H2CO3 content assuming an ideal mixing (Eq. (2)) and the actual H2CO3 content calculated assuming conservation of alkalinity and conservation of DIC (e.g., no degassing, no dissolution, no precipitation). The computed value is 0.7 ± 0.5 mol (Table 7). ii) Dissolution of carbonate minerals (XCO3) according to H2CO3 + XCO3 → X2+ + 2 HCO− 3 (where X is Ca or Mg). Integrating the ΔDIC and ΔDIC⁎ values evaluated above (Table 5) over the pumped volumes, the total amount of added DIC (QDIC and QDIC⁎) is estimated to be between 16.7 and 12.9 mol, respectively. Since the conversion of H2CO3 into HCO− 3 does not change DIC, the mean consumption of H2CO3 corresponds to about 14.8 mol (Table 7). + iii) Cation exchange according to: H2CO3 + Na–Y → HCO− 3 + Na + H–Y (Y is the mineral surface). The maximum amount of H2CO3 potentially consumed is given by the total amount of released Na+ ion (i.e. 11.8 mol; Table 7). iv) Dissolution of minerals other than carbonates, silicates mostly, + converts H2CO3 into HCO− 3 according to H2CO3 + XB → X + − HCO3 + HB (where X is Ca or Mg and B is a residual silicate mineral). The amount of H2CO3 consumed by this process is evaluated by the total dissolved solid cation (TDScat = [Na+] + [K+] + 2 [Ca2+] + 2 [Mg2+]) integrated over the pumped volumes of 35 ± 4 mol and corrected for the Ca2+ charge input due to carbonate dissolution, − 29.6 ± 2 mol. The resulting value is 5 ± 5 mol (Table 7). The molar balance of injected and consumed H2CO3 by these four processes is zero within their error bars. After three weeks of incubation, approximately 60% of the injected H2CO3 has reacted forming HCO− 3 . Less than 2% of this amount occurred via acid–base reactions through mixing between the background water and the injected solution. Roughly 30% of this amount reacted by dissolution of carbonate minerals, that is, by ionic trapping, and the remaining 30% reacted by cation exchange and dissolution of silicate minerals. Using data from this experiment alone, it is difficult to determine the relative contribution of the two latter processes, which both compete to convert H2CO3 into HCO− 3 . 5. Conclusions The methodology developed in this study uses DIC isotope composition and major ion data (DIC, Na+, Ca2+, Mg2+, K+, and ∑Si) to identify processes that convert H2CO3 into HCO− 3 : dissolution of carbonate minerals, cation exchange, and the dissolution of primary minerals such as Ca–Mg silicates during a field-scale CO2 injection test in a reactive formation. In this natural and heterogeneous system, these field results confirm that several geochemical processes lead to the neutralization of the injected CO2. The dominate process was

determined to be carbonate mineral dissolution followed by cation exchange reactions and/or Ca–Mg silicate mineral dissolution in the host formations. Relative contribution were difficult to determine in this setting due to limited precision and accuracy of DIC measurements as well as the spatial and temporal variability of the background water compositions. The role of ionic exchange also remains to be quantified and should be taken into account. For similar applications in the future, quantifying the contribution of these different processes is important for overall understanding and modelling of how CO2 is geochemically trapped in reactive storage reservoirs. Acknowledgments This study was financially supported by the Centre de Recherches sur le Stockage Géologique du CO2, Institut de Physique du Globe de Paris-Total-Schlumberger ADEME partnership. We thank Didier Jezequel for his assistance in experimental analysis, Jean-Paul Toutain (Toulouse University) for his comments on the discussion part, Pierpaolo Zuddas (IPGP) for his participation in fieldwork, W. Masterson and G. Myers (Lamont–Doherty Earth Observatory) for their assistance during fieldwork. We are grateful to Michel Girard (IPGP) for the maintenance of the mass spectrometers and to Nicole Vassard (IPGP) for the help during the preparation of the samples. We would like to thank the two anonymous reviewers and the guest editor Dr Bénézeth Pascale for their constructive comments to improve the whole manuscript. References Andre, L., Audigane, P., Azaroual, M., Menjoz, A., 2007. Numerical modeling of fluid–rock chemical interactions at the supercritical CO2–liquid interface during CO2 injection into a carbonate reservoir, the Dogger aquifer (Paris Basin, France). Energy Convers. Manag. 48, 1782–1797. Assayag, N., Rivé, K., Ader, M., Jézéquel, D., Agrinier, P., 2006. Improved method for isotopic and quantitative analysis of dissolved inorganic carbon in natural water samples. Rapid Commun. Mass Spectrom. 20, 2243–2251. Bachu, S., Adams, J.J., 2003. Sequestration of CO2 in geological media in response to climate change: capacity of deep saline aquifers to sequester CO2 in solution. Energy Convers. Manag. 44, 3151–3175. Bachu, S., Gunter, W.D., Perkins, E.H., 1994. Aquifer disposal of CO2: hydrodynamic and mineral trapping. Energy Convers. Manag. 35, 269–279. Bickle, M., Chadwick, A., Huppert, H.E., Hallworth, M., Lyle, S., 2007. Modelling carbon dioxide accumulation at Sleipner: implications for underground carbon storage. Earth Planet. Sci. Lett. 255, 167–176. Brosse, E., Magnier, C., Vincent, B., 2005. Modeling fluid–rock interaction induced by the percolation of CO2-enriched solutions in core samples: the role of reactive surface area. IFP Revue-Oil Gas Sci. Technol. 60, 287–305. Burgdorff, K., Goldberg, D., 2001. Petrophysical characterization and natural fracturing in an olivine–dolerite aquifer. Electron. Geosci. 6, 1–28. Carroll, S.A., Knauss, K.G., 2005. Dependence of labradorite dissolution kinetics on CO2(aq), Al(aq), and temperature. Chem. Geol. 217, 213–225. Chow, J.C., Watson, J.G., Herzog, A., Benson, S.M., Hidy, G.M., Gunter, W.D., Penkala, S.J., White, C.M., 2003. Separation and capture of CO2 from large stationary sources and sequestration in geological formations. J. Air Waste Manage. 53, 1172–1182. Deines, P., Langmuir, D., Harmon, R.S., 1974. Stable isotope ratios and the existence of a gas phase in the evolution of carbonate groundwaters. Geochim. Cosmochim. Acta 38, 1147–1154. Emberley, S., Hutcheon, I., Shevalier, M., Durocher, K., Mayer, B., Gunter, W.D., Perkins, E.H., 2004. Monitoring of fluid–rock interaction and CO2 storage through produced fluid

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