Iron–clay interactions under hydrothermal conditions: Impact of specific surface area of metallic iron on reaction pathway

Chemical Geology 381 (2014) 194–205

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Iron–clay interactions under hydrothermal conditions: Impact of specific surface area of metallic iron on reaction pathway Franck Bourdelle a,⁎, Laurent Truche a,⁎⁎, Isabella Pignatelli a, Régine Mosser-Ruck a, Catherine Lorgeoux a, Christophe Roszypal a, Nicolas Michau b a b

GeoRessources, UMR 7359, Université de Lorraine, Campus Aiguillettes, 54506 Vandœuvre-lès-Nancy, France Andra, Direction Recherche et Développement/Service Colis et Matériaux, Parc de la Croix Blanche, 1/7 rue Jean Monnet, 92298 Châtenay-Malabry CEDEX, France

a r t i c l e

i n f o

Article history: Received 23 January 2014 Received in revised form 30 April 2014 Accepted 2 May 2014 Available online 21 May 2014 Editor: J. Fein Keywords: Iron metal Clays Serpentine Smectite Redox reaction Hydrogen

a b s t r a c t The long-term evolution of minerals in contact with metallic iron is important in various domains such as the earth sciences, materials science, cosmochemistry or industry. As an illustration, iron–clayey rock interactions are notable issues in the framework of secondary alteration processes in chondrites, or in the evolution of steel canister corrosion in projects for high-level nuclear waste repositories. In these contexts, interactions between the geological environment and metallic iron or engineered structures must be assessed with a high level of precision. Therefore, over the last decade, several experimental studies have focused on metallic iron–clay interactions showing the important relationship between the reaction progress and the iron/clay mass ratio. The present investigation demonstrates that, apart from this mass ratio, the specific surface area of metallic iron has a crucial influence, since it impacts the reaction pathway and the ambient physico-chemical parameters of the medium. For this purpose, two original continuous monitoring experiments were performed to measure pH and pressure in real-time, as well as analyze the gas and solution compositions, by bringing the same mass of (a) iron powder (Siron = 0.07 m2/g) or (b) iron grains (Siron ≈ 0.001 m2/g) into contact with claystone (Callovo–Oxfordian claystone, Bure, France) at 90 °C for 3 months. Using iron powder, i.e. the more reactive cast iron (with a corrosion rate of 0.54 mmol/day for iron powder as against 0.01 mmol/day for iron grains), causes an Fe-enrichment of the clay particles, leading initially to the formation of new phases intermediate between interlayered illite–smectite and iron-rich serpentine, followed by conversion into odinite–greenalite. On the other hand, using iron grains make the clay compositions “kaolinitic” with a noticeable I.C. depletion. Meanwhile, the illite–smectite and quartz fractions of the claystone are destabilized, while the mineral transformations control the pH around 7 (+/- 0.3) and prevent the formation of magnetite, thus contradicting the thermodynamic predictions. The present study, which involves in situ monitoring of pH and H2 production, provides some important keys to obtaining better constraints on reaction mechanisms, kinetics and thermodynamic models, aimed at predicting accurate reaction paths and their long-term consequences. © 2014 Elsevier B.V. All rights reserved.

1. Introduction The long-term evolution of minerals in contact with iron is a key process in various domains such as the earth sciences, materials science, cosmochemistry or industry. Indeed, the understanding of interactions between metallic iron and silicate minerals is of primary importance in several current topics, such as the study of extraterrestrial materials (e.g. Brearley, 2006), or the storage of high-level nuclear ⁎ Correspondence to: F. Bourdelle, GeoRessources, UMR 7359 CNRS, Faculté des Sciences et Techniques, Entrée 3B-Campus des Aiguillettes, Université de Lorraine, 54506 Vandœuvre-lès-Nancy, France. Tel.: +33 3 83 68 47 28; fax: +33 3 83 68 47 01. ⁎⁎ Correspondence to: L. Truche, GeoRessources, UMR 7359 CNRS, Faculté des Sciences et Techniques, Entrée 3B-Campus des Aiguillettes, Université de Lorraine, 54506 Vandœuvre-lès-Nancy, France. Tel.: +33 3 83 68 47 13. E-mail addresses: [email protected] (F. Bourdelle), [email protected] (L. Truche).

http://dx.doi.org/10.1016/j.chemgeo.2014.05.013 0009-2541/© 2014 Elsevier B.V. All rights reserved.

waste (e.g. Gaudin et al., 2013 and references therein), which itself motivates archeological studies of natural analogues (e.g. Michelin et al., 2013). In all these cases, the iron–clay interactions and the resulting mineralogical transformations are intimately dependent on the P–T conditions, on the physico-chemical parameters of the medium and on the nature/crystallochemistry of reacting phases, i.e. their compositions, their crystal properties and their specific surface area. Meteorites are among the few natural objects where silicates coexist with metallic iron, and they are consequently the subject of several studies concerning mineral reactions and alteration (e.g. Krot et al., 1995; Kojima and Tomeoka, 1996; Morlok and Libourel, 2013). In fact, the nature of secondary phases in chondrites and the reaction mechanisms are still poorly understood, while the conditions, location and timing of the secondary alteration processes remain highly controversial. As a result, we need to better constrain the iron–silicate interactions, even at low temperatures. On the other hand, several research programs (OECD-

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NEA, 2012) envisage a combination of natural (claystone, granite or salt formations) and engineered (glass, steel canisters, bentonite, reinforced concrete structures) barriers for radioactive waste confinement. Among these different concepts, one of the considered technical options involves the disposal of vitrified waste emplaced in metal canisters and surrounded by a deep clayey rock formation. In this context, iron–clay interactions will take place as soon as the storage conditions become water-saturated and anoxic (e.g. de Combarieu et al., 2011), which will lead to major mineralogical and geochemical transformations of the host rocks. These two examples highlight that the iron–clay interactions must be carefully assessed in terms of reaction paths and kinetics. In both cases, metallic iron can react with the surrounding materials such as silicates, clay minerals, carbonates and oxides, and, under certain conditions, can form new Fe-bearing phases, such as magnetite, siderite, Fe-smectites, serpentine (such as cronstedtite), chukanovite or chlorite even at moderate temperatures (T b 100 °C; e.g. Lantenois, 2003; Perronnet, 2004; Lantenois et al., 2005; Mosser-Ruck et al., 2010; Jodin-Caumon et al., 2012; Lanson et al., 2012; Pignatelli et al., 2013; Pignatelli et al., submitted for publication). Many studies (e.g. de Combarieu et al., 2007; Pierron, 2011; Rivard, 2011) propose batch tests using the Callovo–Oxfordian claystone (COx) to investigate iron–clayey rock interactions under simulated waste storage conditions (T ~ 90 °C). The Callovo–Oxfordian claystone (Lerouge et al., 2011) was identified as a potential host-rock by Andra (French National Radioactive Waste Management Agency) for the French repository program at Bure (Meuse/Haute-Marne, France). This representative claystone contains quartz, carbonates (calcite and dolomite), clay minerals (illite, illite–smectite mixed layers, chlorite, and kaolinite) as major phases, with minor amounts of feldspars and pyrite (Rousset, 2002; Bregoin, 2003; Claret et al., 2004; Gaucher et al., 2004; Yven et al., 2007). In these hydrothermal experiments, the COx sample was reacted in the presence of iron (provided either in powder form or in foils) and with different solutions. The results mainly indicate that the clay minerals are destabilized and become very reactive, the iron being corroded to varying degrees. New Fe-rich phases are crystallized, while the resulting solutions contain very low contents of iron and high concentration of potassium. Nevertheless, comparisons between these previous studies may be considered speculative. This is because the experimental conditions are very different from one experiment to another, especially the nature of the metallic iron used. Moreover, the experimental monitoring is not always appropriate to constrain the progress of the reaction as a function of time. In fact, in parallel with the characterization of the solid products, it is essential to monitor continuously evolving parameters such as pH, pressure, solution and gas compositions throughout the experiment to define the reaction paths occurring in these diverse iron–clay systems. In this context, and as an extension of previous research, the present study aims to define the role of iron-specific surface area on the evolution of iron–clay systems, by means of continuous monitoring throughout the experiments. The experiments were performed at 90 °C using COx, iron powder or iron grains as starting material, in the presence of saline solution, to simulate the nuclear waste repository conditions. In addition to the analysis of solution and gas composition throughout the course of the reaction, the pH and the H2 production were monitored in situ for the first time in iron–clay experiments, thus giving us important keys to obtaining better constraints on reaction mechanisms, kinetics and thermodynamic models. 2. Materials and methods 2.1. Experiment and starting materials Two series of experiment, denoted A and B, were performed in two separate reactors with different types of metallic iron, with the aim of improving our understanding of the impact of iron-specific surface

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area on the evolution of the system. In the first series (A), 1.2 g of powdered iron (99.9% purity, 40 μm, specific surface area Siron = 0.07 m2/g (± 0.03) as measured by Kr-BET) was placed in contact with 12 g of ground COx claystone (sampled at the Bure site, Meuse/HauteMarne, at a depth of 490 m) and 120 g of saline solution (NaCl 0.0201 mol/kg + CaCl2 0.0029 mol/kg, including a small amount of Br (0.25 mmol·l−1) as a conservative tracer) at 90 °C for 3 months. Therefore, the Fe/COx and liquid/COx mass ratios used in these experiments were 1/10 and 10/1, respectively. The synthetic saline solution is assumed to represent the claystone porewater (Rivard, 2011; protocol details in Guillaume et al., 2003). In the second series (B), the iron was provided as non-spherical grains (with longest axis = 800 μm to 2 mm) to decrease the metal specific surface area by about two orders of magnitude (Siron ranging from 0.95 × 10−3 to 1.6 × 10−3 m2/g, calculated from geometric considerations). The masses of iron, COx and solution were unchanged to preserve the same mass ratios in series A and B. After 3 months at 90 °C, the reactor was opened in a glove box ([O2] b 2 ppmv), saturated with argon to avoid sample oxidation. The solid run-products were dried under Ar inside the glove box. In addition, a blank experiment, without iron metal, was carried out to determine the effect of heating the clay on pH, gas and solution compositions in the first days of the run. For this purpose, 12 g of COx was placed in contact with 120 ml of saline solution for 28 days at 90 °C under the same conditions used for A and B experiments. The two reactors used are stainless steel autoclaves lined with Teflon®, equipped with pH and P–T probes for in situ measurements (Fig. 1a). The pressure is estimated with an accuracy of 0.1 bar, while the temperature is obtained with a precision of 0.1 °C. A zirconia pH electrode (Ag/Ag2O) and an Ag/AgCl reference electrode (0.1 mmol·l−1 KCl electrolyte) were supplied by Corr Instruments LLC. These electrodes have wetted metallic parts made of pure titanium grade 3, and can operate at temperature ranging from 60 °C to 200 °C and at pressures up to 140 bars. The e.m.f. developed between the pH and the reference probes is measured with a digital pH meter having an input impedance higher than 1012 Ω. The zirconia pH electrodes and Ag/AgCl electrode pair are equilibrated in heated buffers at 90 °C, and calibrated with a three-point procedure for pH (standards pH25°C = 4 [C6H4(COOH) COOK], 7 [Na2HPO4/KH2PO4] and 9 [Na2B4O7]), before and after each experiment, to check the measurement deviation (Fig. 1b). Therefore, all pH values given in this study are at 90 °C and are corrected for probe deviation. The measured e.m.f. exhibits high stability over long periods of time, with a drift of less than 0.1 pH units between the start and end of experiments (90 days). Calibration curves show a sub-Nernstian slope close to the theoretical value at 90 °C (i.e., 72.01 mV/pH). Before the beginning of the run, the autoclaves were flushed and bubbled with an argon flow (2–5 bars) during 30 min to ensure anoxic conditions, and were sealed with a starting pressure of ~ 1.1–1.3 bars of argon (measured at 25 °C). Gas and solution sampling was carried out throughout the run using sealed and flushed storage cells made of stainless steel, by means of two valves connected to the reactor (Fig. 1a). No sampling was carried out in the reactor dedicated to pH measurements, so the data could be obtained in an undisturbed system. 2.2. Characterization techniques Elemental analysis of sampled solutions was performed by inductively coupled plasma optical emission spectrometry (ICP-OES) for Na, Ca, Si, K, S, Al, Fe and Mg after acidification in HNO3 (dilution × 10; precision of ±10%), and ionic chromatography (IC) for carbonates, Cl− and Br− (dilution in milliQ water 3 times; precision of ±2%). The gas samples were analyzed by gas chromatography (GC) coupled to a thermal conductivity detector (TCD), previously calibrated using several Ar + H2 + CO2 mixtures of different concentrations at various pressures injected from an injection loop of 20 μl. The column is a Carbobond (Agilent, 50 m in length, internal diameter of 0.53 mm,

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(EDX) fitted to the microscope, with a counting time of 40 s and a dead time lower than 15%. The TEM K factors (Cliff and Lorimer, 1975) were calibrated on standards, with a maximum error of 5% for each element. Given the high spatial resolution, analysis points were targeted on isolated clay particles, to reveal compositional variations of the finest fraction of the run-product. For each sample, 30 to 50 particles were analyzed to be representative. The analysis consists of measuring the atomic% of Si, Ti, Al, Fe, Mg, Na, Ca and K. The structural formulae are calculated on the basis of O10(OH)2, assuming that the interlayer cations are Na, Ca, K, and that all iron is ferrous. The TEM specimens were prepared by dispersing the b 2 μm fraction of run-product powders in ethanol under ultrasonication, and evaporating a drop of suspension on a network-like holey carbon support film placed on a 200 mesh copper grid. A Scanning Electron Microprobe (SEM) images were also acquired using a FEG Hitachi S-4800 instrument, operated at 15 kV with a current of 1.5 nA, to examine the morphology of iron grains before and after the experiment. 2.3. Thermodynamic calculations The thermodynamic data presented in this study are calculated using PHREEQC geochemical software package (Parkhurst and Appelo, 1999) associated with the THERMODDEM database (http://thermoddem. brgm.fr; Gaucher et al., 2009; Blanc et al., 2012), based on the LLNL database (Johnson et al., 2000). PHREEQC geochemical simulations are used to calculate the solubility of minerals at 90 °C, as well as Eh–pH stability fields of the Fe–H2O system at 90 °C and carbonate-system equilibrium parameters, such as Ca2+ content, carbonate content and CO2 partial pressure (PCO2). 3. Results 3.1. In situ measurements and gas analysis

Fig. 1. (a) Schematic diagram of the experimental set-up used for continuous monitoring of iron–clay experiments; and (b) in situ pH probe deviation over time after regular recalibration at 90 °C. According to this set-up, the pH is measured and corrected for deviation during the experiment, along with checking of the temperature and pressure.

film thickness of 10 μm), traversed by helium as carrier gas (constant pressure of 1.03 bar), allowing precise identification of Ar, CO2 and H2 gases. The temperature program consists of two steps of column heating, with a first step at 30 °C during 2 min, and a second step at 180 °C during 5 min, using a heating rate of 50 °C/min. The solids were characterized as oriented and disoriented powders by X-ray diffraction (XRD), using a D8 Bruker diffractometer with non-monochromatic CoKα radiation and suitable analytical conditions (35 kV accelerating voltage, 45 mA current, a step size of 0.035 or 0.02 °2θ and a counting time of 3 s or 4 s per step over a 2–60 or 4–23 °2θ range depending on the sample). Air-dried and ethylene glycol (EG)saturated oriented samples were prepared from the b 2 μm fractions of the run-products to identify clay minerals. A Transmission Electron Microscope (TEM) study was carried out with a Philips CM200 instrument equipped with an Si–Li detector, operated at 200 kV, to obtain semi-quantitative chemical analyses, as well as to examine the morphology of newly formed crystals. The chemical analyses were obtained using an energy dispersive X-ray analyzer

Fig. 2a reports the in situ pH values obtained from experiments conducted at 90 °C in dedicated reactors. In experiment A, where the COx sample is in contact with powdered iron, the time-evolution of pH displays three distinct consecutive stages: (i) an initial decrease from 7.61 down to 7.03 after 3 days, corresponding probably to equilibration of the system with the COx mineralogy, (ii) a slight increase during 24 days, reaching a maximum value of 7.2, and then (iii) a slow decrease to a constant value of 6.89 after 90 days. On the other hand, experiment B with iron grains shows a slightly different evolution of pH, which decreases during the first 16 days down to 6.78, followed by a slight increase to reach a value of 6.88 after 50 days, and then slowly decreases again to reach 6.8 after 90 days elapsed time (Fig. 2a). Thus, experiment B shows a pH evolution similar to the blank experiment during the first few days; the pH in the blank experiment is almost completely stabilized after 25 days at around 6.71, as predicted by Tournassat et al. (2008) and Truche (2009). While the temperature is maintained at 90 °C in the system, the pressure is free to change during the course of the reaction. The pressure increase due to emission of different gases was monitored precisely, and the nature of these gases sampled during the experiments is characterized by chromatographic analysis. The results summarized in Table 1 indicate that the excess of pressure mainly corresponds to H2 release. In the experiment with iron powder (A), the pressure of H2 is almost 4 bars at the 90th day; in the experiment with iron grains (B), the H2 content increases continuously during the experiment, to reach 0.35 bar at the end of the run. The CO2 released represents a very low fraction of the gases present in the case of experiment A (11 mbar), whereas it decreases from 26 mbar at the 18th day to 16 mbar at the 75th day of experiment B. According to these results, after correction for PCO2, PAr, vapor pressure (0.7 bar at 90 °C) and pressure losses due to gas sampling, the in situ continuous pressure measurement may be taken as representing the evolution in H2 content (Fig. 2b). On the

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basis of this pressure monitoring, H2 content rapidly increases during experiment A, from 0 at the beginning of the run to 3.8–4.2 bars at the 50th day, after which it starts to stabilize to reach a constant value of 4.3 after 60 days. Thus, hydrogen is rapidly released in large quantities in the presence of iron powder, while it is slowly released in the system with iron grains. In the case of the blank experiment, the pressure excess measured after 12 and 27 days is very weak (b 38 mbar), and corresponds exclusively to CO2 degassing due to the attainment of carbonate equilibrium (Table 1). 3.2. Solution analysis The sampled solution compositions are reported in Fig. 3. In experiment A (COx + powdered iron), K concentration increases significantly and continuously from 1.37 to 3 mmol·l−1 during 90 days, while Mg content first decreases during the first 30 days, from 0.69 (Mg stored in the COx and released into the medium) to ~ 0.05 mmol·l− 1, and then remains constant probably controlled by the dolomite solubility. Ca concentration ranges between 2.21 and 2.87 mmol·l− 1, i.e., near the concentration of the starting saline solution (2.9 mmol·l−1). Na content is also maintained around the value of the starting solution (20.7 mmol·l−1), and Si content increases slightly up to 0.6 mmol·l−1 during the run, but remains under-saturated with respect to quartz at 90 °C. Fe and carbonate concentrations are lower than 0.03 mmol·l−1 and 0.9 mmol·l−1, respectively, and S content is between 0.6 and 1.2 mmol·l−1. The measurement accuracy does not allow us to compare the Fe concentration with the magnetite and siderite solubilities, which are close to 10−7 mmol·l−1 and 0.01 mmol·l−1 at 90 °C, respectively. Concerning the tracers, the concentrations of Cl and Br remain constant, near the initial value (25.6 and 0.25 mmol·l−1, respectively), confirming the conservation of mass in solution, and the low water consumption induced by iron oxidation. In the case of experiment B (COx + iron grains), the Mg concentration slowly decreases, from 0.95 mmol·l− 1 (18th day) to 0.66 mmol·l− 1 at the end of the run, i.e. 10 times higher than the value obtained in experiment A (COx + iron powder). K concentration also increases during the run, from 1.09 to 1.65 mmol·l−1. On the other hand, concentrations of Na and Si vary similarly as in experiment A. Ca and S remain around 3.8 and 1.6 mmol·l−1, respectively, while Fe is close to the detection limit, except for two measurements that yield anomalously high concentrations probably due to a sample contamination. Only the Cl tracer was added to the solution of this experiment, and its concentration is stable throughout the run at around 30 mmol·l−1. Carbonate contents could not be measured. Analysis of the blank experiment solution shows the same cationic content variations as those observed in the first thirty days of the B series experiments, starting from 0.01 mmol·l−1 for Fe, 1 mmol·l− 1 for K, 2.9 mmol·l− 1 for Ca, 0.7 mmol·l− 1 for Mg, and 0.1 mmol·l−1 for Si at t = 0. Therefore, the evolution of cation and anion concentrations is very similar during the A, B and blank experiments (except for Mg), albeit with less pronounced variations in the blank and B experiments, indicating very slight mineral transformations. 3.3. XRD characterization of solid run-products Fig. 4a shows XRD patterns recorded on disoriented preparation of A and B experiment run-products (bulk rock). These patterns are compared to those obtained from the starting material (unreacted COx + metallic iron) and two short additional experiments of 2 and 4 Fig. 2. (a) pH and (b) H2 partial pressure variation vs. time, during A, B and blank experiments. Each curve is plotted from over 60 analytical points (20 for blank experiment). Since H2 degassing accounts for over 98% of the pressure excess (see text and chromatographic results), the variation of pressure presented here can be taken as PH2 variation. Pressure measurements are compared with the analyses of gases sampled regularly over the course of the experiment. The discontinuities in the pressure curves (in gray) correspond to liquid and/or gas sampling. H2 degasing rate is presented in inset (c).

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Table 1 Gas chromatography analyses performed during the course of experiments A and B, expressed in molar fraction and partial pressure, recalculated from initial pressure, initial PAr, and vapor pressure at 90 °C. A-pH: reactor dedicated to pH measurements. *PCO2 partial pressure of CO2 expected from PHREEQC tests, obtained by fixing solution compositions and pH values at different sampling times (see text) and assuming calcite equilibrium. Experiment

Days

A A A-pH B B B B B Blank Blank

12 90 90 18 36 54 75 90 12 27

Molar fraction (%mol)

Partial pressure (bar)

H2

Ar

CO2

H2

Ar

CO2

49.78 91.76 78.54 11.11 16.43 20.14 26.99 31.88 0.0 0.0

50.21 8.00 21.22 87.23 81.88 78.40 71.66 65.78 97.25 96.57

0 0.25 0.24 1.66 1.70 1.46 1.36 2.34 2.75 3.43

1.09 (±0.2) 4.04 (±0.2) 3.93 (±0.2) 0.17 (±0.2) 0.23 (±0.2) 0.26 (±0.2) 0.32 (±0.2) 0.35 (±0.2) 0.0 0.0

1.10 (±0.1) 0.35 (±0.1) 1.06 (±0.1) 1.34 (±0.1) 1.14 (±0.1) 0.99 (±0.1) 0.85 (±0.1) 0.72 (±0.1) 1.07 (±0.1) 1.06 (±0.1)

0.000 (±0.02) 0.011 (±0.02) 0.012 (±0.02) 0.026 (±0.02) 0.024 (±0.02) 0.019 (±0.02) 0.016 (±0.02) 0.026 (±0.02) 0.030 (±0.02) 0.038 (±0.02)

days in which powdered iron was in contact with COx at 90 °C. The results indicate that quartz is partially dissolved during experiment A (decreasing intensity of the 24.5 °2θ CoKα reflection in run-products sampled at 2, 4 to 90 days), while metallic iron (52.5 °2θ CoKα peak) is totally consumed after 90 days. On the other hand, the mineralogical assemblage of experiment B stays very similar to that of the starting material. Although iron grains are still present in the run-products of

PCO2 (bar) expected⁎ 0.004 0.013 0.013 0.021 0.016 0.013 0.012 0.010 – –

experiment B, iron peaks are not observed on the XRD pattern. This result is due to the fact that, because of their large size, iron grains were not included in the preparation of the sample for XRD analysis. The contents of chlorite, calcite, dolomite, feldspar and kaolinite appear unmodified from XRD results, and the chlorite and kaolinite peaks appear more clearly defined after heating, showing a better crystallinity. Finally, the presence of newly-formed iron oxides such as magnetite is

Fig. 3. Time-evolution of solution compositions during experiments A and B, showing total dissolved content of Ca, Fe, and K, Mg, Si and carbonates. The solubilities of quartz and dolomite in saline solution at 90 °C are highlighted. The error bars indicate assumed analytical precision, i.e., ±10%.

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Fig. 4. XRD patterns of starting material and solid run-products: (a) disoriented powder of bulk rock, (b) oriented powder of extracted fine fraction (b2 μm), with background subtracted. EG: ethylene-glycol-saturated samples; AD: air-dried samples. Indicated phases: C: chlorite; I: illite/muscovite; I/S: illite–smectite mixed layers; 7 Å: kaolinite and/or serpentine-like phases, Qz: quartz; Ca: calcite; D: dolomite/ankerite; F: feldspar; P: pyrite, Fe: iron powder. The bulk-rock XRD patterns (a) are normalized to the calcite peak. XRD patterns are shifted vertically for clarity (intensity in arbitrary units).

not revealed by XRD, and serpentine-like phases are not clearly revealed because of the superimposition of their peaks with chlorite and kaolinite. For the b2 μm fraction of the samples, a similar comparison shows that the typical reflections of interstratified clays observed on the AD and EG patterns of the starting material decrease in intensity or disappear in experiments B and A, respectively (Fig. 4b). The major reflection between 7.5 and 9 °2θ CoKα typical of interstratified clays shifts toward lower angles on XRD patterns of the AD treated samples of the A and B experiment run-products, probably representing a partial dissolution of the illite component of I/S mixed layer mineral during the experiments, and a correlated enrichment of smectite in the sampled clay particles. The XRD patterns obtained after ethylene glycol (EG) saturation of the samples confirm a rapid dissolution of I/S in the case of experiment A, but their continued existence in experiment B even after 90 days (Fig. 4b). 3.4. TEM and SEM investigations The results of TEM–EDX analyses on the b2 μm fraction of runproducts from experiments A and B are plotted in Fig. 5 using three

different ternary projections: a 4Si–M+–R2 + diagram (R2 + refers to divalent cations) with clay mineral end-members (Meunier and Velde, 1989), an Mg–AlVI–Fe diagram (octahedral composition) and an Na–2Ca–K diagram (interlayer charge). In these projections, the COx starting material plots in a field intermediate between illite, celadonite, high-charge beidellite and high-charge montmorillonite, which is comparable to illite and mixed-layer illite/smectite clay compositions. In addition, several analyses plot near to the kaolinite composition. When compared to the starting material, clay particles from experiment A are enriched in Fe and slightly depleted in octahedral Al. The Mg and Si contents are close to those measured in the initial COx, while the content of interlayer cations greatly decreases (I.C. = Na + 2Ca + K lower than 20% in 4Si–M+–R2+ diagram), mainly because of K depletion. Ca content does not vary significantly, whereas Na content increases from nearly 0 at.% in the clay particles of the COx to 10–40 at.% as interlayer cations in experiment A particles. Indeed, the clay particles evolve during the reaction toward a serpentine composition, plotting across a field ranging from mixed-layer illite/smectite to Fe-saponite or Fe-serpentine/chlorite.

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intermediaries between 7 Å Fe-rich phase and smectite, as observed by Mosser-Ruck et al. (2010), but which are poorly crystallized. Except for several hairy particles, the clay particles show little change in terms of crystallinity, crystal-size or crystal-shape, and the chemical composition therefore remains the most discriminant character. Magnetite is not observed, and amorphous gels are not identified. SEM observations of iron grains after reaction (experiment B) were carried out on cut-polished sections of the coarse fraction embedded in resin, showing partial and very limited dissolution, mainly observed along metal fractures (Fig. 6d). Iron grains are well preserved in terms of shape and size. At the SEM scale, the identification of a gel is not possible. Although a layer of newly formed Fe-rich silicates can be observed on the surface of iron grains (Fig. 6e), it is not possible to determine at this scale if this new phases could cause grain passivation. In the case of experiment A, no residual iron powder could be found by SEM and TEM examination, in agreement with XRD results and indicating the complete corrosion of iron during the reaction. Therefore, by comparison between the two experiment series, the iron corrosion rate seems to be controlled by the reaction surface. 4. Discussion 4.1. Time-evolution of the powdered iron–clay system: pH evolution and reaction progress

Fig. 5. Chemical compositions of solid run-products and initial material obtained by TEM– EDX analysis, plotted in ternary diagrams: 4Si–M+–R2+ diagram with end-members (Meunier and Velde, 1989), Mg–AlVI–Fe diagram (octahedral composition) and Na–2Ca– K diagram (interlayer charge). Be: beidellite; Mon: montmorillonite; Mus: Muscovite; Phe: phengite; Cel: celadonite; Sap: saponite; Stev: stevensite; Bio: biotite; Serp: serpentine; Chl: chlorite. The gray arrows indicate the evolution of clay particle composition from the COx starting material in both experiments A and B, the question mark refers to the hypothetical character of the marked evolution.

Clay particles of experiment B undergo less modification than those of experiment A, i.e., their chemical compositions are closer to the COx starting material, especially the octahedral sheet. This is probably related to the smaller specific surface of the iron grains, limiting their reactivity with COx claystone and/or the quantity of Fe2 + released. Therefore, the particles from experiment B contain less Fe but more Al compared with particles from experiment A. The clay particles are also characterized by (i) a major loss of K, (ii) a slight gain of Si, similar to experiment A, and (iii) a less pronounced Na enrichment, with Ca remaining as the predominant interlayer cation. In summary, the composition of particles from experiment B plots in the illite–smectite–kaolinite field, with a noticeable I.C. depletion. The morphology of the b 2 μm fraction particles was extensively investigated and example pictures are reported in Fig. 6. No significant morphological variations are detected when comparing particles from experiments A and B with the COx as the starting material; the typical flake aspect of clay minerals is preserved and dominant in the run products (Fig. 6a). TEM observations also show that particles having the composition of Fe-rich T–O phyllosilicates, such as greenalite, are not observed in experiments A or B. Only some iron-rich particles from experiment A show a newly formed hairy morphology (Fig. 6b). These kinds of particles may correspond to berthierine-like phases or

Most of the clay particles (except for chlorite) resulting from experiment A are of poor crystallinity, but their composition appears rather homogeneous. Therefore, in terms of crystal chemistry, newly-formed particles range from interstratified illite–smectite (I/S) to Fe-rich silicates, either serpentine or chlorite, while the b2 μm fraction of the starting claystone is representative of pure illite and illite–smectite. According to the XRD results and the evolution of K concentration in solution, the initial I/S seem to be rapidly destabilized, in particular the illite layers of the I/S (Jodin-Caumon et al., 2012), while the iron released into solution by the corrosion of iron powder is immediately incorporated into new clay mineral phases. These particles are depleted in K, but still contain Na–Ca-smectite layers and have a composition far from iron-rich serpentine. Because of the unfavorable experimental conditions used in this study, in particular the low iron/solid ratio and/or the duration of the experiment, the crystallization of pure serpentine phases is probably limited. Indeed, several authors, using slightly different conditions but employing the same starting claystone, observe the formation of pure 7 Å-phases. Using liquid/solid mass ratios of 20 and 10, and iron powder/COx mass ratios of 0.2 and 0.1 or 1, Pierron (2011) and Rivard (2011) report the formation of odinite and greenalite after a 9-month experiment at 90 °C (Fig. 7). Using a liquid/solid mass ratio of 10 and an iron/COx ratio of 0.5, Pignatelli et al. (2013) obtain a better development of iron-rich clays by decreasing the experimental temperature step by step from 90 °C to 40 °C, identifying these phases as cronstedtite. Fig. 7 shows a Fe–Si–Al diagram comparing the present results with those of Pignatelli et al. (2013), indicating that the reaction path modifies the 10 Å-phase composition toward Fe-phases via odinite and greenalite, but not toward berthierine. Thus, the particles observed here reflect the first step of conversion from 10 Å- to 7 Å-phases. The other phases of the COx, such as calcite or feldspars, seem to be in equilibrium with the system, while the quartz is partially dissolved, due to the under-saturation of the solution in Si with respect to this phase. The precipitation of new iron-rich silicates is an active process that controls Si concentration at a low level in the bulk solution. The non-clay fraction of the COx, apart from quartz, is not really reactive, but contributes to stabilizing the pH between 6.9 and 7.2, while the addition of Fe0 should lead to a major alkaline perturbation, according to the iron corrosion equation: 0

2þ

Fe þ 2H2 O ¼ Fe

−

þ 2OH þ H2 :

ð1Þ

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Fig. 6. TEM (a, b, c; b2 μm fraction) and SEM (d, e) images of the solid run-products in experiments A and B: (a) Flake morphology of poorly crystallized A particles; (b) hairy morphology of A particles with the highest Fe content; (c) flake morphology of B particles; (d) slightly corroded iron grains in experiment B; and (e) iron grain rim covered by a layer of Fe-rich silicates.

A large increase in pH could be expected in the A series experiments because of the complete and rapid consumption of iron. Similarly, Bildstein et al. (2006) and Marty et al. (2010) have predicted a pH value stabilized at around 10–11 under various conditions using thermodynamic models simulating long-term iron–clay interactions. Rivard et al. (2013) measured a pH value of 9 in an iron-liquid experiment, i.e., in the absence of clays. In the same way, using a PHREEQC simulation of iron dissolution in saline solution with a solid/liquid mass ratio of 0.1, but without mineral phases, we can predict an increase of the pH up to 9.8. Surprisingly, such high values are not observed here, possibly because the claystone system has an unexpected capacity to control the pH. Therefore, after reaching a value of 7 (5th day), the pH starts to increase slightly as a consequence of the iron corrosion, but the induced mineralogical transformations produced in the COx claystone keep the pH under control. Finally, the pH value stabilizes at around 6.9 after the complete dissolution of iron. However, the competition between iron corrosion and rock alteration for the control of acidity in the medium depends not only on the nature of the starting

material, in terms of the specific surface area, rock mineralogy (claystone, pure smectite or bentonite) and mass ratios but also on the nature of the corrosion products. At 80 °C, Lantenois et al. (2005) recorded pH values of around 9.5 starting with pure smectite, whereas de Combarieu et al. (2007) measured pH values up to 10.05 with a high iron/COx ratio and a low solid/liquid ratio. In our experiments, the excess of COx and the precipitation of Fe-rich phyllosilicates maintain the pH close to neutrality, the mineralogical transformations being fast enough to attenuate and even neutralize the pH increase induced by the iron oxidation. Therefore, the iron content remains very low in solution, and the released Fe2+ is immediately incorporated into the newly-formed Fe-rich phyllosilicates. The influence of the carbonate fraction on the pH control capacity of claystone is determined through PHREEQC simulation tests; the partial pressures of CO2 calculated at each sampling time are reported in Table 1. Fixing the starting solution composition ([Ca2 +] = 2.9 mmol·l−1, [Na+] = 20.1 mmol·l−1 and [Cl−] = 25.9 mmol·l−1) and a pH value of 6.9, corresponding to the final pH measured in A series

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Fig. 7. Chemical compositions of solid run-products and initial material, obtained by TEM– EDX analysis, plotted in a Fe–Si–Al ternary diagram. The compositional analyses are compared to those obtained for well-developed newly formed phases by Pierron (2011), Rivard (2011) and Pignatelli et al. (2013). End-members are: Gr: greenalite (Guggenheim et al., 1982); Od: odinite (Bailey, 1988); Ber: berthierine (Brindley, 1982); and Cr: cronstedtite (Kogure et al., 2002). The reaction path followed by Fe powder–clay interactions is highlighted by a gray area.

experiments (Fig. 2), the thermodynamic calculation of solution-calcite equilibrium yields 0.013 bar for PCO2, 0.87 mmol·l−1 for carbonate concentration in solution and 3.9 mmol·l−1 for Ca2+ content. These calculated results are in agreement with the experimental observations (Fig. 3 and Table 1), even though the Ca2 + content is slightly overestimated by 1 mmol·l−1. When fixing PCO2 (0.011 bar as measured at the end of experiment A, Table 1) instead of pH, the PHREEQC test yields 6.94 for the pH value, 0.78 mmol·l−1 for carbonate concentration in solution and 3.7 mmol·l−1 for Ca2+ content, these values being similar to those obtained in the first test. In the same way, the carbonate system can be simulated for the 12th day of the experiment, fixing the pH at 7.15: the results are 0.004 bar for PCO2, 0.49 mmol·l−1 for carbonate concentration in solution and 3.4 mmol·l−1 for Ca2+ content, i.e., very close to the experimental data (Table 1 and Fig. 3). This cross-checking shows that (i) the carbonate system is equilibrated in the medium at each sampling time and for each pH, and (ii) both the Ca2+ content in solution and CO2 partial pressure are constrained by calcite solubility at 90 °C. Therefore, the role of the calcite system is subordinate to the bulk rock effect, and contributes along with this latter to the stabilization of the medium around a neutral pH. 4.2. Time-evolution of the iron grain-clay system: role of specific surface area of iron particles When iron is added to the experiment as grains (B case), the clay particles of the solid run-products are less evolved. From XRD, solution

analysis and TEM investigations, a small amount of the starting I/S remains after 3 months of reaction. As in the case of experiment B, the newly-formed phases are depleted in K and enriched in Na, but are also Fe-poor, possibly due to the formation of kaolinite-like minerals even though no typical polygonal kaolinite was observed by TEM. In the present study, it is difficult to confirm the crystallization of these new clay minerals, because (i) most of the kaolinite–smectite particles have smectite-like morphology (Cuadros et al., 2009) and (ii) the analyzed particles have a mixed saponite–kaolinite composition rather than a mixed smectite–kaolinite composition (Fig. 5). The population of particles analyzed in experiment B has the same composition in a Fe–Si–Al diagram than the starting COx and no iron-rich serpentinelike minerals are reported (Fig. 7). Although the solids do not display noticeable iron enrichment (Fig. 5), the iron concentration in solution remains very low (close to the detection limit). The lack of Fe release is thus due to the nature of the iron metal introduced into the system: the specific surface area of iron grains is very low, which slows down the corrosion process. A passivation layer at the surface of the iron grains could also explain the very low Fe2 + content in solution, but the SEM data do not confirm this hypothesis. During the first two weeks of the experimental run, the pH is controlled by the claystone rather than by iron dissolution, and reaches a value of 6.78 (see comparison between B and blank experiments, Fig. 2a). Then, when I/S are altered and the Fe2 + content increases in the system, the pH rises slowly until the end of the run. As in the case of experiment A, the carbonate system behavior is simulated with PHREEQC, using the B series experimental parameters. By fixing the pH value at 6.78 (Fig. 2), the first simulation test yields 0.021 bar for PCO2, 1.13 mmol·l− 1 for the carbonate content in solution and 4.08 mmol·l− 1 for the Ca2 + content. In the second test, where PCO2 is fixed at 0.026 bar (as measured, Table 1), we obtain 6.73 for the pH value, 1.3 mmol·l−1 for the carbonate content in solution and 4.2 mmol·l−1 for the Ca2+ content. These values are similar to those measured in situ and confirm the equilibrium of the carbonate system with the medium and the dependence of the pH on the mineral assemblage. In conclusion, a comparison between experiments A and B highlights the fact that the amount of corroded iron is higher with iron powder than with iron grains, modifying the nature of the solid runproducts. Although the comparison of the present results with previous studies shows the strong impact of iron/clay ratio and T conditions on the progress of conversion of 10 Å to 7 Å-phases, the role played by the specific surface area of metallic iron also appears to be significant: (i) the formation of serpentine-like phases rather than “kaolinitic” particles (arrows in Fig. 5) is dependent on the synchronization between the mineralogical transformations and iron release, which is subordinated to the specific surface area of iron particles; (ii) the evolution of pH is driven by both the mineralogical transformations and the iron corrosion process, which depends on the specific surface area of the iron (Fig. 2a); (iii) the corrosion rate of the iron depends on its specific surface area, thus highlighting a change in the corrosion process (see Section 4.3.2). Therefore, the specific surface area of the iron reactant is an additional key parameter along with the iron/clay ratio, which controls not only the kinetics but also the reaction path of iron–clay interactions. 4.3. Iron corrosion 4.3.1. Magnetite precipitation Metallic iron in contact with clays often oxidizes to magnetite, as observed in many studies (e.g., de Combarieu et al., 2007; Gaudin et al., 2009; Schlegel et al., 2010; Jodin-Caumon et al., 2012; Rivard et al., 2013). However, magnetite is not identified in the present study

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Fig. 8. Simplified Eh–pH diagram for iron-bearing phases, considering Fe2+, Fe3+, hematite, magnetite, greenalite and stability field of water, calculated at 90 °C with PHREEQC software, for a concentration of 0.1 × 10−5 mol·l−1 (according to the average Fe2+ concentration measured in experiment A solutions). The gray-dashed lines represent the measured pH window and the gray-shaded area the plausible Eh range, in accordance with experimental results (see text).

as a newly-formed phase. At the measured Fe2+ concentrations, the pH values required to cross the magnetite stability field should be higher than 7.3 (Fig. 8), according to the simplified Eh–pH diagram of iron presented in Fig. 8 (considering Fe2+, Fe3+, hematite, magnetite, greenalite and stability field of water) calculated at 90 °C with PHREEQC software. This pH condition is never reached in our experiments. In fact, the pH and the Fe2+ concentrations are buffered at low levels by COx alteration, the rapid formation of Fe-rich phases in experiment A and the slow dissolution of iron grains in experiment B (Fig. 8). Moreover, the degassing of H2, the anoxic conditions and the absence of hematite provide evidence of a very low negative Eh value. These combined factors therefore prevent the formation of magnetite (Fig. 8). This conclusion is similar to that proposed by Gaudin et al. (2013), who used thermodynamic simulations of the interactions between iron and Tournemire claystone to identify a range of conditions (especially the buffering of pore-waters at near neutral pH by the claystone) at which magnetite is absent. In fact, the iron/clay mass ratio is not the only parameter that plays a role in the formation of magnetite. 4.3.2. Monitoring of H2 degassing as a precise probe for the measurement of iron dissolution rate The hydrogen pressure increase should be controlled by the iron corrosion rate (Eq. (1)). Assuming that the amount of H2 adsorbed at the surface of clays is negligible in the present conditions (Didier et al., 2012) and that H2 solubility is governed by Henry's law, the amount of H2 produced during experiment A can be estimated at 19.8 mmol (16.3 mmol of H2(g) from maximum pressure value + 3.5 mmol of H2(aq)), whereas the amount of H2 theoretically produced by the total oxidation of Fe0 initially present in the system into Fe (II) (Eq. (1)) is 21 mmol. The small discrepancy between the theoretical and the measured amount of H2 produced (b6%) is within the analytical uncertainty and demonstrates that Fe0 mostly oxidized to Fe2 + and that the degassed H2 can be considered as a conservative reaction product, i.e., chemically inert. Indeed, the potential reduction of pyrite into pyrrhotite is hindered by the low H2 partial pressure and the neutral pH, which are unfavorable conditions for pyrrhotite precipitation

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(Truche et al., 2013). As a result, H2 production is an excellent probe for measuring iron corrosion rate. Therefore, considering the initial linear parts of the H2 pressure variation curves (Fig. 2c), we can estimate the iron corrosion rate as being equivalent to H2 production rate over the first twenty days, which gives values of 0.54 mmol·day−1 and 0.011 mmol·day−1 for experiments A and B, respectively. Assuming that the specific surface area of iron remains constant in the first stage of the reaction, the corrosion rates can be expressed as mmol/day/m2, varying from 4.5 to 11.25 for iron powder and from 5.7 to 9.65 for iron grains, according to the ranges of values for a specific surface area. These mean corrosion rates are very high, and in the same order of magnitude for both iron powder and iron grains. In their study, de Combarieu et al. (2007) using an iron/clay mass ratio of 1 (Siron = 0.26 m2/g) calculate a mean corrosion rate of around 0.5 mmol/day/m2, which is much lower than the values proposed here. The values reported by de Combarieu et al. (2007) are derived from XRD patterns, and represent indirect and less precise estimations, whereas the present study proposes a direct measurement of mean corrosion rate through the monitoring of H2 production. However, apart from the differences in methodology of these two corrosion rate studies, a comparison of the results shows an apparent negative correlation between the corrosion rate (in mmol/day/m2) and the specific surface area of iron. This phenomenon was previously observed by Liu et al. (2008) studying the dissolution of sedimentary and hydrothermal pyrites. By comparing pyrite dissolution rate and the specific surface area of grains measured by BET, these authors point out that the expected linear relationship between these two parameters is not systematically observed. According to these authors, the discrepancy can be explained by the pH variations and the differences in morphology defects (e.g. McKibben and Barnes, 1986) and crystal surface development (e.g. Kendelewicz et al., 2004), as well as the porosity and purity of particles. In the present experiments, the iron grains are nonspherical highly angular particles (as Fig. 6d), which can enhance the iron dissolution rate by the preferential alteration of corners, cracks and micro-pores of grains. All these considerations imply that the iron/clay ratio, and consequently the mineral assemblage of the COx impact the corrosion process in conjunction with the specific surface area of the iron particles, thus modifying or controlling the physico-chemical characteristics of the system. 5. Conclusion This study shows the complexity of iron–clay interactions, demonstrating that the specific surface area of metallic iron has a major effect on the reaction path beyond the iron/clay mass ratio. In fact, the nature of the newly-formed phases (“kaolinitic” or Fe-rich particles) depends on whether or not the iron corrosion/Fe2+ release processes are synchronous with the alteration of the claystone (quartz, illite). The specific surface area of the iron mainly influences the reaction pathway (from 10 Å-phases to kaolinitic or Fe-rich particles), while the iron/clay and liquid/clay ratios mostly impact the reaction rate, i.e., the development of new phases either smectitic serpentine or pure serpentine. In this context, using an iron/clay mass ratio of 0.1, the present experiments show the rapid dissolution of powdered iron and the crystallization of smectite-Fe-rich silicates formed from the dissolution of I/S (mostly the illite fraction), which represents a first step in the transformation to odinite–greenalite. On the other hand, the reaction of clays with iron grains leads to the crystallization of a K-depleted phase with preserved Al content. In both cases, the pH seems to be control around the neutrality by the iron–claystone interaction rather than by the iron corrosion alone, while the latter is mainly responsible for the large pressure increase induced by H2 degassing. From these observations, we infer that the continuous monitoring of pH and PH2 during experimental runs is a crucial approach to better understand the influence of each parameter on the progress of the reaction. Moreover, the

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measurement of pH at T and P during the course of the experiment allows us to show that the alteration of COx strongly impacts the time-evolution of pH (avoiding high pH values) and prevents the precipitation of magnetite. Since H2 may be considered as a conservative reaction product resulting only from iron corrosion, it could become a valuable probe for measuring the iron corrosion rate in the case of an abiogenic scenario. Due to differences between laboratory experiments and in situ conditions, these results are insufficient on their own to allow direct extrapolation of the conclusions to a waste repository or meteorite alteration. Nevertheless, the present study gives important keys for obtaining better constraints on thermodynamic/kinetic models, with the aim of predicting reaction paths as well as their consequences to a very high level of confidence. Acknowledgments Gilles Bessaques and Aurélien Randi are warmly thanked for their technical assistance. The authors thank L. Salsi for SEM images acquired at the SCMEM laboratory (Univeristé de Lorraine, France), S. Migot for TEM images and analyses at IJL (Institut Jean Lamour, Centre de Compétences en Microscopies Electroniques et Microsondes (CC-MEM)), and Geraldine Kitzinger, Hervé Marmier and David Billet for the analysis of solutions carried out at LIMOS laboratory (Université de Lorraine, France). We would like to thank Dr. Lietai Yang (Corr Instrument LLC) for constructive discussions and continued improvement of the pH probes used in this study. This research was financially supported by Andra (Agence Nationale pour la gestion des Déchets Radioactifs / French National Agency for the Management of Radioactive Wastes) I11ABC3ANDRA. The discussions and comments of the journal editor Jeremy Fein and of two anonymous reviewers are gratefully acknowledged. The authors also thank Michael Carpenter for his professional English language review. References Bailey, S.W., 1988. Odinite, a new dioctahedral–trioctahedral Fe+ 3 -rich 1:1 clay mineral. Clay Miner. 23, 237–247. Bildstein, O., Trotignon, L., Perronnet, M., Jullien, M., 2006. Modelling iron–clay interactions in deep geological disposal conditions. Phys. Chem. Earth, Part A/B/C. 31, 618–625. Blanc, P., Lassin, A., Piantone, P., Azaroual, M., Jacquemet, N., Fabbri, A., Gaucher, E.C., 2012. Thermoddem: a geochemical database focused on low temperature water/rock interactions and waste materials. Appl. Geochem. 27, 2107–2116. Brearley, A.J., 2006. The Action of Water in Meteorites and Early Solar System. In: Lauretta, D.S., McSween, H.Y. (Eds.), Arizona University Press, pp. 587–624. Bregoin, S., 2003. Variabilité spatiale et temporelle des caractéristiques du CallovoOxfordien de Meuse/Haute-Marne. (Paris, Thesis) , p. 258. Brindley, G.W., 1982. Chemical compositions of berthierines — a review. Clay Clay Miner. 30, 153–155. Claret, F., Sakharov, B.A., Drits, V.A., Velde, B., Meunier, A., Griffault, L., Lanson, B., 2004. Clay minerals in the Meuse–Haute marne underground laboratory (France): possible influence of organic matter on clay mineral evolution. Clay Clay Miner. 52, 515–532. Cliff, G., Lorimer, G.W., 1975. The quantitative analysis of thin specimens. J. Microsc. 103, 203–207. Cuadros, J., Nieto, F., Wing-Dudek, T., 2009. Crystal–chemical changes of mixed-layer kaolinite–smectite with progressive kaolinization, as investigated by TEM-AEM and HRTEM. Clay Clay Miner. 57, 742–750. de Combarieu, G., Barboux, P., Minet, Y., 2007. Iron corrosion in Callovo–Oxfordian argilite: from experiments to thermodynamic/kinetic modelling. Phys. Chem. Earth 32, 346–358. de Combarieu, G., Schlegel, M.L., Neff, D., Foy, E., Vantelon, D., Barboux, P., Gin, S., 2011. Glass–iron–clay interactions in a radioactive waste geological disposal: an integrated laboratory-scale experiment. Appl. Geochem. 26, 65–79. Didier, M., Leone, L., Greneche, J.M., Giffaut, E., Charlet, L., 2012. Absorption of hydrogen gas and redox processes in clays. Environ. Sci. Technol. 46, 3574–3579. Gaucher, E., Robelin, C., Matray, J.M., Négrel, G., Gros, Y., Heitz, J.F., Vinsot, A., Rebours, H., Cassagnabère, A., Bouchet, A., 2004. Andra underground research laboratory: interpretation of the mineralogical and geochemical data acquired in the Callovian– Oxfordian formation by investigative drilling. Phys. Chem. Earth Part A/B/C 29, 55–77. Gaucher, E., Tournassat, C., Pearson, F.J., Blanc, P., Crouzet, C., Lerouge, C., Altmann, S., 2009. A robust model for pore-water chemistry of clayrock. Geochim. Cosmochim. Acta 73, 6470–6487. Gaudin, A., Gaboreau, S., Tinseau, E., Bartier, D., Petit, S., Grauby, O., Foct, F., Beaufort, D., 2009. Mineralogical reactions in the Tournemire argilite after in-situ interaction with steels. Appl. Clay Sci. 43, 196–207.

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