Embed Size (px)
Citation preview
Please print only if necessary – Please use double-sided printing – Please recycle paper – Thanks!
Please print only if necessary – Please use double-sided printing – Please recycle paper – Thanks!
Weathering of olivine under CO2 atmosphere: A martian perspective 1
2
E. Dehouck a,*, A. Gaudin a, N. Mangold a, L. Lajaunie b, A. Dauzères c, O. Grauby d, E. Le 3
Menn a 4
5
aLaboratoire de Planétologie et Géodynamique de Nantes (LPGN), CNRS/Université de 6
Nantes, 44322 Nantes, France 7
bInstitut des Matériaux Jean Rouxel (IMN), CNRS/Université de Nantes, 44322 Nantes, 8
France 9
cInstitut de Radioprotection et de Sûreté Nucléaire (IRSN), Fontenay-aux-Roses, France 10
dCentre Interdisciplinaire de Nanoscience de Marseille (CINaM), CNRS/Aix-Marseille 11
Université, Campus de Luminy, 13288 Marseille, France 12
13
14
*Corresponding author. Now at: Department of Geosciences, State University of New York at 15
Stony Brook, Stony Brook, NY 11794-2100, USA. 16
E-mail address: [email protected] 17
18
19
2
ABSTRACT 20
21
Recent analyses from the Curiosity rover at Yellowknife Bay (Gale crater, Mars) show 22
sedimentary rocks deposited in a lacustrine environment and containing smectite clays 23
thought to derive from the alteration of olivine. However, little is known about the weathering 24
processes of olivine under early martian conditions, and about the stability of smectite clays in 25
particular. Here, we present a 3-month experiment investigating the weathering of forsteritic 26
olivine powders (Fo90) under a dense CO2 atmosphere, and under present-day terrestrial 27
conditions for comparison. The experiment also evaluates the potential effects of hydrogen 28
peroxide (H2O2), as a representation of the highly oxidizing compounds produced by 29
photochemical reactions throughout martian history. The weathered samples were 30
characterized by means of near-infrared spectroscopy (NIR), X-ray diffraction (XRD), 31
transmission electron microscopy with energy dispersive X-ray spectrometry (TEM-EDX), 32
Mössbauer spectroscopy and thermogravimetry. The results show that a Mg-rich smectite 33
phase formed from the weathering of olivine under CO2 conditions, although in lower 34
abundance than under terrestrial conditions. The main secondary phase formed under CO2 35
turns out to be a silica-rich phase (possibly acting as a “passivating” layer) with a non-36
diagnostic near-infrared spectral signature. The use of H2O2 highlights the critical importance 37
of both the redox conditions and Fe content of the initial olivine on the nature of the 38
secondary phases. 39
40
Keywords: Mars, alteration, olivine, smectites, phyllosilicates, silica. 41
42
43
3
1. INTRODUCTION 44
45
The martian crust is known to have undergone aqueous alteration in its early history: 46
various secondary minerals – including clays (mostly Fe-Mg smectites), hydrated sulfates, Fe-47
(oxy)hydroxides and carbonates – have been identified during the last decade by 48
hyperspectral imagery, almost exclusively in terrains of Noachian or Hesperian age (Bibring 49
et al., 2005; 2006; Poulet et al., 2005; Carter et al., 2013). While the detections of these 50
secondary minerals and the involvement of water in their formation are widely accepted, the 51
climatic and geological settings in which such processes occurred remain controversial. The 52
correlation between hydrous minerals and valley networks, alluvial/delta fans and paleolakes 53
in ancient terrains (e.g., Ehlmann et al., 2008a; Dehouck et al., 2010; Ansan et al., 2011) 54
suggest that at least some of the secondary minerals were formed by rock-water-atmosphere 55
interactions, such as weathering. Numerous observations of vertical sections of Al-rich clays 56
over Fe/Mg-smectites, typical of pedogenesis processes, also favors such a scenario (e.g., 57
Loizeau et al., 2007; 2010; Ehmann et al., 2009; Murchie et al., 2009; Wray et al., 2009; Noe 58
Dobrea et al., 2010; Gaudin et al., 2011; Le Deit et al. 2012; Carter et al., 2013). However, 59
this hypothesis has been challenged based on the “missing” widespread carbonates – which 60
are expected to form in a thick CO2 atmosphere – (e.g., Bullock and Moore, 2007; Fernandez-61
Remolar et al., 2011) and the difficulty to model an early martian atmosphere having the 62
adequate properties for liquid water to be available (e.g., Forget et al., 2012; Wordsworth et 63
al., 2012). This led some authors to propose that most of the martian clays have been formed 64
at depth (i.e., isolated from the atmosphere) by hydrothermal groundwater circulation, and 65
later exposed at the surface by erosion (e.g., Ehlmann et al., 2011). 66
Therefore, it is important to better understand how primary rocks and minerals 67
weather under Mars-relevant surface conditions in order to evaluate if the secondary 68
4
mineralogy currently observed on Mars may derive from rock-water-atmosphere interactions 69
(i.e., weathering), or if other processes (deep, hydrothermal alteration) should be favored 70
instead. Specifically, it is important to clarify if Fe/Mg-smectites – which are the most 71
common clay minerals found on Mars (Carter et al., 2013) – can be produced by weathering 72
under a dense CO2 atmosphere and associated low pH, because it is known that low pH 73
conditions tend to destabilize Fe/Mg smectites (e.g., Galan et al., 1999; Bauer et al., 2001; 74
Murakami et al., 2004; Amram and Ganor, 2005; Golubev et al., 2006) and to favor Al-rich 75
clays instead (e.g., Komadel and Madejová, 2006; and references therein). Yet, X-ray 76
diffraction analyses from the Curiosity rover of fine-grained lacustrine sedimentary rocks 77
display patterns consistent with a proportion of ~20% of a smectite clay, most probably a 78
ferroan saponite, suggesting that significant alteration occurred at Gale crater (Vaniman et al., 79
2014). The lack of high temperature phases (e.g., serpentine, chlorite) and the overall 80
mineralogy suggest a formation by alteration of forsteritic olivine through low-grade 81
diagenesis (McLennan et al., 2014), thus highlighting the role of olivine in alteration 82
processes at this place. 83
In this context, only few experimental studies so far have compared the weathering of 84
primary silicates under terrestrial and simulated martian conditions. Accordingly, the main 85
goal of this study was to determine the effects of early martian conditions on the weathering 86
of forsteritic olivine (Fo90), with a particular focus on the secondary products. Owing to its 87
nesosilicate structure (unpolymerized silica tetrahedra), olivine is one of the most reactive 88
silicates in near-surface conditions and, for this reason, is particularly appropriate for such 89
study. Using four batch experiments, we evaluated the effects of a CO2 atmosphere, but also 90
those of highly oxidizing compounds (represented here by hydrogen peroxide, H2O2) that 91
have been detected on present-day Mars and may have been more abundant in the past. A 92
detailed chemical, petrographic and mineralogical characterization of the final samples was 93
5
achieved using an appropriate set of analytical techniques, including near-infrared 94
spectroscopy (NIR), X-ray diffraction (XRD), transmission electron microscopy with energy 95
dispersive X-ray spectrometry (TEM-EDX), Mössbauer spectroscopy and thermogravimetry. 96
Implications for the alteration processes on early Mars are discussed. 97
98
2. APPROACH AND METHODS 99
100
2.1. Background: previous work on olivine alteration under CO2 101
102
Because of its high chemical reactivity, olivine has been frequently used in laboratory 103
experiments examining the dissolution rates of silicates in various conditions (e.g., Wogelius 104
and Walther, 1992; Pokrovsky and Schott, 2000; Golubev et al., 2005; Hänchen et al., 2006). 105
Olsen and Rimstidt (2007) have summarized the results of numerous of these studies and used 106
them to evaluate the lifetime of an olivine grain exposed to weathering on the martian surface. 107
Taking several parameters into account (temperature, composition, grain size, pH), they 108
estimated values ranging from a few thousand and several million years (up to 30 Ma for a 109
forsterite grain of 1 mm at 273 K and a pH of 7.5). 110
A number of experiments have also examined the potential of forsteritic olivine for 111
CO2 sequestration, in the context of Earth’s climate change mitigation (e.g., Giammar et al., 112
2005; Gerdemann et al., 2007; Bearat et al., 2006; Garcia et al., 2010; King et al., 2010; Daval 113
et al., 2011). These experiments have led to a good understanding of the olivine carbonation 114
process. The most common secondary phases reported are magnesite and amorphous silica. 115
Although the results of these studies are not directly applicable to weathering processes on 116
early Mars because of their high temperatures (up to several hundred degrees C) and CO2 117
pressures (up to several hundred bars), they can provide relevant information about the 118
6
mechanisms and environmental parameters that control the alteration of olivine under CO2 119
conditions. Nevertheless, the discrepancy between the results of such studies and the lack of 120
widespread carbonates on Mars justifies studying the alteration of olivine in conditions that 121
are more relevant to surface processes in terms of temperature and pressure. 122
Lastly, the experiments conducted by Dehouck et al. (2012) included the weathering 123
of forsteritic olivine, among other silicates, under simulated early martian conditions – 0.8 bar 124
of CO2, with or without H2O2 – and at low liquid-to-rock (L/R) ratio. The authors reported the 125
formation of nesquehonite, an hydrated Mg-carbonate. 126
127
2.2. Selection of the starting material 128
129
As a starting material, we selected a natural sample of forsteritic (Mg-rich) olivine 130
from San Carlos, Arizona (USA). Olivine is a common component of ultramafic and mafic 131
rocks, from which the martian crust is mainly made. It has been unambiguously detected in 132
various regions of Mars based on orbital data (Hoefen et al., 2003; Mustard et al., 2005; 133
Poulet et al., 2007; Ody et al., 2013) and, despite the global enrichment in Fe of the martian 134
crust (e.g., Longhi et al., 1992), most of the detections correspond to forsteritic and/or fine-135
grained olivine (both Fe content and grain size have a similar influence on the NIR signature 136
of olivine, so that it is difficult to retrieve definitive compositions from orbit; Ody et al., 137
2013). Moreover, among the secondary minerals discovered so far in olivine-rich regions of 138
Mars (Mangold et al., 2007; Ehlmann et al., 2009; Gaudin et al., 2011; Bishop et al., 2013), 139
several species of phyllosilicates (smectites, serpentines) and carbonates (magnesite, 140
hydromagnesite) are known to derive from olivine or olivine-bearing rocks in terrestrial 141
environments (e.g., Delvigne et al., 1979; Velbel et al., 1991; Wilson, 2004). Finally, 142
forsteritic olivine is weakly resistant to weathering, but it is also sufficiently common on the 143
7
surface of the Earth – as opposed to fayalitic olivine, in particular – to be found in “fresh” 144
state and in sufficient amount (>150 g) to meet the requirements of this study. 145
146
2.3. Weathering conditions on early Mars 147
148
Designing a laboratory experiment under “simulated early martian conditions” 149
requires making some assumptions. Based on the present-day martian atmosphere, composed 150
of >95% CO2 (Owen, 1982), the easiest way to obtain higher surface temperatures is by 151
increasing the total pressure. Recent results from general circulation models suggest that 152
additional greenhouse gases or other factors (impacts, volcanism) may have been necessary to 153
reach temperatures above the freezing point (e.g., Forget et al., 2012). However, since these 154
factors are still poorly constrained, we have decided to consider only the effect of a dense CO2 155
atmosphere in the present study. Other experiments are devoted to test the effect of SO2 on the 156
weathering processes (e.g., Chevrier et al., 2012). 157
Another hypothesis of our study is that the highly oxidizing compounds found in the 158
regolith (Bullock et al., 1994; Zent, 1998; Yen et al., 2000; Hurowitz et al., 2007) and in the 159
atmosphere (Clancy et al., 2004; Encrenaz et al., 2004) of present-day Mars may have been 160
more abundant at the time of formation of the alteration minerals. Indeed, these compounds 161
derive from the UV-induced photolysis of water; thus, if water was more abundant on early 162
Mars, highly oxidizing compounds are expected to be more abundant, too. In our experiment, 163
these oxidizing compounds have been represented by hydrogen peroxide (H2O2). Although 164
limited to the first meters or first tens of meters of the soil (Bullock et al., 1994; Zent, 1998), 165
hydrogen peroxide combined with a dense CO2 atmosphere may have influenced alteration 166
processes over long time scales by allowing the existence of both (moderately) acid and 167
(highly) oxidizing solutions (e.g., Chevrier et al., 2006; Hurowitz et al., 2010; Fernandez-168
8
Remolar et al., 2011). Yet, only few experimental studies to date have explored its effect on 169
the weathering of silicates (Dehouck et al., 2012). Because our experimental device is 170
exposed to sunlight and heat, H2O2 is expected to rapidly undergo disproportionation into 171
H2O and O2 during the course of the experiment. However, since disproportionation does not 172
involve transfer of an electron, the redox state of the system is not changed and the effect of 173
highly oxidizing conditions can still be tested. 174
175
2.4. Protocols 176
177
2.4.1. Starting material: preparation and characterization 178
179
The forsteritic olivine used in this study originated from San Carlos, Arizona (USA) 180
and consisted of olive-green, centimeter-sized monocrystals. Preliminary SEM-EDX 181
(scanning electron microscopy coupled with energy dispersive X-ray spectrometry) analyses 182
(N=24) performed on polished grains confirmed the mean formula to be ~Fo90Fa10 (Fo: 183
forsterite; Fa: fayalite). They also revealed the presence in minor quantity of small (a few tens 184
of micrometers for the longest ones; Fig. 1), elongated crystals containing P, Ca as well as a 185
little F, and interpreted to be apatite. 186
The cleanest grains were selected from the original lot and minor bright impurities 187
macroscopically visible on the surface of some grains were removed using a diamond-covered 188
abrasion tool. The grains were then cleaned using an ultrasonic bath in ethyl alcohol and 189
finely crushed using an automatic crusher with agate balls. The resulting, white-colored 190
powder was homogenized. SEM observations showed that >98% of the powder grains were 191
<60 µm in size. 192
9
To monitor the purity of the olivine powder before the experiment, we performed a 193
precise mineralogical and chemical characterization using near-infrared spectroscopy (NIR), 194
powder X-ray diffraction (XRD), and inductively coupled plasma optical emission 195
spectroscopy (ICP-OES). The NIR spectrum showed a highly pure olivine with very limited 196
hydration (Fig. 1). Shallow absorption bands at 2.31-µm (Hunt et al., 1972) and 4.02 µm are 197
attributable to minor apatite (Fig. 1). XRD pattern showed no crystallized phase other than 198
olivine. Finally, the chemical analysis by ICP-OES (Table 1) again confirmed the 199
composition of the olivine to be Fo90Fa10. Apart from Si, Mg and Fe, other elements like Mn, 200
Ca, Al and Na are also detected in small amounts. In contrast, P is below the detection limit, 201
which allows determining an upper limit of 0.12 wt% for the quantity of apatite contained in 202
the powder. 203
204
2.4.2. Experimental device 205
206
For the purposes of this study, we designed and assembled the apparatus shown in Fig. 207
2. Its primary function is to provide controlled weathering conditions in terms of atmospheric 208
composition and temperature. It consists of four borosilicate glass (Schott® Duran) flasks – or 209
“reactors” hereafter – of 1 liter, each placed on a hot plate. The first two reactors are 210
hermetically closed and linked to a stainless-steel and PTFE circuit allowing introduction of 211
gaseous CO2 (Air Liquide® Alphagaz CO2 N45). Independent gas lines and valves ensure 212
that each reactor remains fully isolated from the second one. A pair of manometers provides a 213
direct visual control of the internal pressure of the two reactors. The apparatus also allows 214
sampling both the solution and the gas without exposing the samples to ambient air. Lastly, a 215
stainless-steel thermometer is immersed inside the solution to continuously control its 216
temperature. 217
10
The two other reactors are filled with ambient air, but are also hermetically closed to 218
prevent escape of water vapor or other gases originating from the solution. They are not 219
equipped for pressure monitoring or gas sampling. 220
221
2.4.3. Initial protocol 222
223
For each reactor, we precisely weighted 30 g of sample and poured it into 300 mL of 224
liquid, thus giving a L/R ratio of 10. In two reactors, the liquid consisted only of ultrapure 225
water with a resistivity of 18 MΩ.cm. In the two other reactors, it consisted of hydrogen 226
peroxide (H2O2; VWR® Prolabo GPR Rectapur) at ~1 vol% in ultrapure water. 227
The four reactors were then mounted on the experimental device. The two reactors 228
used for the terrestrial conditions (one with H2O2 – hereafter abbreviated as “Air-H2O2” – and 229
the other without H2O2 – hereafter “Air”; Fig. 2) were simply closed with their original 230
stopper and put on the hot plates. The two reactors used for simulated early martian 231
conditions were connected to the circuit described above, which includes its own stoppers. 232
Then, the two reactors were purged by a continued flow of CO2 (the solutions “bubbled”), 233
while the valve for gas sampling stayed opened (Fig. 2). The purge was maintained during 234
one hour for the reactor without H2O2 (hereafter abbreviated as “CO2”) in order to completely 235
remove dissolved nitrogen and oxygen. In contrast, this operation was limited to a few 236
minutes for the other reactor containing H2O2 (hereafter “CO2-H2O2”) since hydrogen 237
peroxide was expected to produce O2 anyway. The apparatus was then closed and the initial 238
CO2 pressure in the “Mars reactors” was set at ~1.5 bar. This value slightly above the ambient 239
atmospheric pressure was chosen to ensure that any tiny leak would cause the escape of CO2 240
and not the entrance of ambient air. This value is also well within the range of atmospheric 241
11
pressure usually explored by global circulation models of early Mars (e.g., Forget et al., 242
2012). 243
Lastly, the temperature of the hot plates (outside the reactor) was set at ~70 °C in 244
order to obtain a solution temperature of ~45 °C (± 5 °C, depending on the room 245
temperature). Aluminum sheets were wrapped around the reactors to help maintaining a 246
homogeneous temperature. The value of 45 °C was chosen as a compromise between two 247
opposite requirements: (1) increasing reaction rates to maintain the experiment in a reasonable 248
timeframe and (2) keeping realistic conditions for surface aqueous processes. To avoid a non-249
homogeneous distribution of the secondary products within the olivine powders due to the 250
slight temperature gradient within the solution, the reactors were manually shaken daily 251
during the course of the experiment. 252
253
2.4.4. Monitoring, solution sampling and final samples retrieval 254
255
Throughout the experiment, a daily reading of temperature and pressure was achieved, 256
so that adjustments could be made if necessary. In particular, our apparatus would allow the 257
reintroduction of gaseous CO2 in the “Mars reactors” in the case of a pressure decrease under 258
a security level that we placed at 1.2 bar. However, this operation never became necessary, 259
demonstrating the efficient airtightness of the apparatus (our estimated leak rate is <1 mbar 260
per day). 261
As expected, hydrogen peroxide rapidly underwent disproportionation into H2O and 262
O2 under the combined action of sunlight and heat. As a result, the pressure inside the “CO2-263
H2O2” reactor reached 2.8 bar after 3 days (1.5 bar of CO2 and 1.3 bar of O2). The same 264
process occurred inside the “Air-H2O2” reactor, with an expected – but not verified – total 265
pressure of 2.3 bar (1 bar of air and 1.3 bar of additional O2). 266
12
The solutions of the four reactors were sampled after 3, 15 and 31 days. A final 267
sampling was made after 95 days, just before the end of the experiment. At each sampling, 268
~10 mL of solution was pumped using a syringe, so that the decrease of the L/R ratio after 30 269
days was relatively limited (from 10 to 9). Each aliquot was filtered at 0.2 µm and its pH was 270
immediately measured. Note here that for both “Earth reactors”, the pH values at day 3 were 271
dismissed because of unstable measures (probably due to the low concentrations of 272
electrolytes in the solutions; see section 3.1.2). Then, the solution was acidified at 2% with 273
nitric acid (HNO3) and analyzed for major elements – Na, Mg, Al, Si, P, K, Ca, Ti, Mn, Fe – 274
by ICP-OES at the SARM-CRPG laboratory (Vandoeuvre-lès-Nancy, France). A blank 275
sample (obtained from a 3-month blank experiment using ultrapure water but no solid) was 276
also analyzed in order to check any possible contamination originating from the experimental 277
device or from the equipments used to manipulate the sampled solutions. 278
Concerning the “Air-H2O2” reactor, the opening of the stopper caused the escape of 279
excess O2 originating from the disproportionation of hydrogen peroxide. As a consequence, 280
the appropriate quantity of liquid H2O2 was poured inside the reactor after each sampling in 281
order to replace the lost O2. 282
The solution sampling at 95 days marked the end of the experiment. The hot plates 283
were turned off and a gas sampling was achieved for the two “Mars reactors” (see next 284
section). Two distinct methods were used to collect the weathered olivine powders: the “CO2” 285
and “CO2-H2O2” reactors were vacuum-dried in order to avoid contact with the room 286
atmosphere, whereas the “Air” and “Air-H2O2” flasks were dried in an oven at 50°C. In both 287
cases, a portion of the solution was first extracted with a syringe to facilitate the drying 288
process. The drying procedure lasted ~2-3 days for each reactor. 289
290
2.4.5. Gas sampling and analysis 291
13
292
Our experimental device is equipped with a gas sampling outlet for each of the “Mars 293
reactors” (Fig. 2). This sampling was achieved only at the end of the experiment (95 days) to 294
avoid any disturbance of the system. The internal air of the stainless-steel bulbs used for the 295
sampling was pumped for one hour before allowing the gas of the reactors to enter. Then, the 296
samples were analyzed by gas chromatography (GC) at the Subatech laboratory (École des 297
Mines, University of Nantes, France). To calibrate the instrument, we used the CO2 bottle of 298
the experimental device (Air Liquide® Alphagaz CO2 N45) as well as two gas mixtures 299
optimized for low O2 concentrations (0.5% O2–99.5% CO2 and 5% O2–95% CO2; Air 300
Liquide® Crystal). Outside this range and for any other gases, measured values must be 301
considered as semi-quantitative. 302
As expected, the analysis of the gas from the “CO2-H2O2” reactor shows a high 303
quantity of O2 produced by the disproportionation of hydrogen peroxide (~2/3 of the mixture, 304
but because of the calibration procedure, the uncertainty is large). In addition, a very small 305
quantity of dihydrogen (H2) is detected (<0.1%). 306
The measured O2 concentration in the “CO2” reactor is minimal: 0.17%, i.e. in the 307
same order of magnitude as in the present-day atmosphere of Mars (Owen, 1982). 308
Furthermore, this value is an upper limit, since some small contamination may have occurred 309
during the preparation and analysis of the gas sample. H2 is detected at a higher level than for 310
the “CO2-H2O2” reactor (~0.5%, so higher than O2), indicating that an oxidation process 311
occurred in the absence of free O2, by dissociation of H2O molecules (e.g., Hurowitz et al., 312
2010). In contrast, no traces of methane have been found (Neubeck et al., 2011). 313
314
2.5. Analytical methods for the solid samples 315
316
14
All analytical methods described below were applied to the bulk powder samples, and 317
some of them were also applied to the <2-µm fraction of these powders, which was separated 318
by settling according to Stokes’ law and retrieved by centrifugation. This was done in an 319
attempt to facilitate the detection and identification of potential clay minerals, which are 320
concentrated in the finer fraction when present as individual particles. Also, smaller olivine 321
grains have higher surface-to-volume ratio, which makes secondary phases easier to detect 322
relatively to coarser fractions. 323
NIR spectra of our initial and weathered powders were acquired at the LPGN 324
laboratory (Nantes, France) using a Nicolet® 5700 Fourier transform infrared spectrometer 325
(FTIR) equipped with a tungsten-halogen white-light source, a CaF2 beam splitter and a 326
DTGS detector. For each sample, the FTIR chamber was first purged for four minutes with 327
dry and CO2-free air. Spectra presented in this paper are the average of 200 measurements in 328
the wavelength range of 1-5 µm (10000-2000 cm-1) with a resolution of 4 cm-1. Background 329
spectra were acquired using a Labsphere® Infragold reference, with a rough surface 330
optimized for powder analyses. Data acquisition and background correction were done using 331
the OMNIC® software. NIR spectra were acquired for both bulk samples and <2-µm 332
fractions. Unfortunately, the <2-µm fraction of the initial olivine was polluted during 333
preparation due to the failure of an ultrasonic agitator, making it unusable for NIR 334
spectroscopy. 335
Powder X-ray diffraction patterns were acquired at the IMN laboratory (Nantes, 336
France) using a Siemens® D5000 diffractometer equipped with a copper source (operated at a 337
voltage of 40 kV and a current of 40 mA), a monochromator and a Moxtek® PF2400 Si PIN 338
detector. For bulk powders, measurements were done between 2θ = 3.5° and 70° with steps of 339
0.016° and a counting time per step of 150 s. Oriented mounts were prepared from the <2-µm 340
fractions by wet smearing on glass slides in order to facilitate the detection of (001) 341
15
diffraction peaks of potential clay minerals. Measurements of the oriented mounts were done 342
between 2θ = 3.5° and 33° with steps of 0.016° and a counting time per step of 320 s. 343
Transmission electron microscopy (TEM) observations were performed at the IMN 344
laboratory using a Hitachi® HF 2000 equipped with a cold-field emission gun operated at an 345
acceleration voltage of 100 kV. TEM-EDX analyses were performed at the CINaM laboratory 346
(Marseille, France) using a JEOL® JEM 2010 equipped with a LaB6 electron gun (operated at 347
200 kV) and a Bruker® EDX detector. 348
57Fe Mössbauer spectra were acquired at the Mössbauer facility of the IMN laboratory, 349
which is equipped with a room-temperature 57Co(Rh) source. The spectrometer is operated in 350
transmission geometry and in constant acceleration mode, with a symmetric velocity ranging 351
between ±4.5 mm/s. Calibration of the velocity scale was done using absorption lines of iron 352
foil. After folding, the spectra (256 channels) were computed with a least-squares routine 353
using Lorentzian lines. 354
Thermogravimetric analyses (TGA) were performed at the CEA laboratory (Saclay, 355
France) using a Netzsch® STA 409 PC LUXX thermobalance operated at a heating rate of 356
10 °C/min from 25 to 980 °C, under a dinitrogen flow of 60 mL/min. Data were then 357
converted into differential thermogravimetric (DTG) curves. 358
Lastly, measures of total inorganic carbon (TIC) were performed at the IRSN 359
laboratory (Fontenay-aux-Roses, France) using an Elementar® Vario TOC Cube. This 360
instrument has a limit of detection of the order of ppb. 361
362
3. RESULTS 363
364
3.1. Chemistry of the solutions 365
366
16
3.1.1. pH evolution 367
368
Measurements show that the pH is more acidic in the two reactors containing a CO2 369
atmosphere than in those containing terrestrial air. In the “Mars reactors”, the pH reaches a 370
value of ~6 after 3 days and remains nearly unchanged afterwards, whereas in both “Earth 371
reactors” it reaches 8.4 at the end of the experiment (Fig. 3). The comparison of these 372
measures with theoretical initial values of 3.9 for a CO2 atmosphere (PCO2=1) and of 5.6 for a 373
terrestrial atmosphere (PCO2=4.10-4) indicates a strong buffering effect of olivine crystals, 374
especially fast in the “Mars reactors”. This suggests a rapid consumption of protons due to the 375
hydrolysis of olivine. 376
The pH values for the “CO2-H2O2” and “CO2” reactors are always very close to each 377
other (respectively 6.1 and 6.2 at the end of the experiment). The situation is a bit more 378
complex concerning the “Earth reactors”: the pH values at 14 and 31 days are slightly lower 379
in the “Air-H2O2” reactor than in the “Air” reactor. However, final values are found to be 380
identical (pH=8.4; Fig. 3). Thus, the pH appears unaffected by the presence of H2O2. 381
382
3.1.2. Dissolution of major elements 383
384
Mass concentrations over time of the three most abundant cations of the system (Si, 385
Mg, Fe) are shown in Fig. 3. The complete dataset (10 elements) with associated uncertainties 386
is available in the electronic annex (Table S1). 387
The evolutions of pH and major elements concentrations as a function of time and 388
atmospheric composition are well correlated: (1) as for pH, there is a rapid increase in 389
concentrations in the first days of the experiment followed by a relative stabilization 390
afterwards; (2) concentrations of dissolved elements are clearly much more important in the 391
17
“Mars reactors” than in the “Earth reactors” – acidic conditions under CO2 favor the 392
hydrolysis of olivine and thus the release of elements in the solution; (3) no systematic 393
influence of H2O2 is noticed (although some differences do exist for Mg, see below). 394
At the end of the experiment, Si is about 14 times more abundant in the “Mars 395
reactors” than in the “Earth reactors”. Beyond the values by themselves, the evolution over 396
time is also different between the two types of conditions: under CO2, the concentration is 397
nearly equal at 14 and 31 days, and has slightly decreased at 95 days; in contrast, under air, a 398
slow but continuous increase is observed. The effect of hydrogen peroxide on the dissolution 399
of Si appears negligible, since the two curves (with and without H2O2) are very close to each 400
other for the two types of conditions. 401
Mg is by far the most abundant element in the solutions. At the end of the experiment, 402
its mass concentration is 2.8, 4.1, 3.7 and 1.9 times higher than for Si in the “CO2-H2O2”, 403
“CO2”, “Air-H2O2” and “Air” reactors, respectively. Hence, the latter is the only one which 404
could tend toward a congruent dissolution (corresponding to a ratio of 1.6). However, ratios 405
for the sampling at 3, 14 and 31 days are all above 3. Also, contrary to Si, the curves appear 406
different with or without hydrogen peroxide: the Mg concentration decreases after 31 days in 407
the “CO2-H2O2” reactor, whereas a nearly linear increase is observed after 14 days in the 408
“CO2” reactor. Under terrestrial atmosphere, Mg concentrations increase continuously all 409
along the experiment (Fig. 3). Lastly, while the Mg concentration is higher with H2O2 for the 410
“Mars reactors”, the contrary is observed for the “Earth reactors”. 411
Iron solubility depends on its oxidation state: ferrous iron (Fe2+) is soluble whereas 412
ferric iron (Fe3+) is insoluble. Therefore, it is not surprising that dissolved Fe is solely 413
detected in the “CO2” reactor, which is the only one to have reducing conditions (no O2 or 414
H2O2). The decrease of concentration after 14 days suggests that some Fe2+ initially dissolved 415
has been subsequently incorporated in a Fe2+- or Fe3+-bearing solid phase. Indeed, even 416
18
without free O2 available, oxidation is possible through the dissociation of H2O molecules, as 417
shown by the detection of H2 in the GC analyses. 418
After Si, Mg and Fe, the most abundant elements are K, Ca and Na. The latter is the 419
only one to have a higher concentration under terrestrial atmosphere than under CO2 (Table 420
S1). Finally, the other analyzed elements (Al, P, Ti and Mn) present very low to non-421
measurable concentrations. 422
423
3.2. Weathered olivine powders 424
425
3.2.1. Near-infrared spectroscopy 426
427
The overall shape of the NIR spectra obtained from our weathered samples (Fig. 4) is 428
still dominated by the spectral signature of olivine (especially the broad, Fe2+-related band at 429
1.04 µm), which is not surprising given the relatively limited timeframe and low temperature 430
of the experiment. However, some significant differences can be observed between 1.8 and 431
2.5 µm (Fig. 4B). In this wavelength range, the four weathered samples show an obvious 432
absorption band at 1.91 µm, which was nearly-absent in the spectrum of the initial olivine. 433
This band is due to the formation of at least one alteration phase containing H2O molecules in 434
its structure (e.g., Hunt, 1977). Furthermore, another absorption band is visible at 2.31 µm in 435
the spectra of the “Air” and “Air-H2O2” samples, but also in the one of the “CO2” sample 436
(although more subtle here). This band indicates the formation of Mg-OH bonds in the altered 437
material (Clark et al., 1990). 438
These visual observations can be confirmed and refined by calculating the band depths 439
for several series of spectra (Fig. 5). Firstly, the band at 1.91 µm turns out to be deeper for the 440
powders weathered under CO2 than for those weathered under terrestrial air. Secondly, the 441
19
band depth at 2.31 µm is indeed greater for the “Air”, “Air-H2O2” and “CO2” samples 442
compared to the initial olivine. Thirdly, for a given type of atmosphere (CO2 or air), the band 443
depth – both at 1.91 and 2.31 µm – tends to be slightly lower for the samples weathered in the 444
presence of H2O2. 445
The combination of a H2O-related band at 1.9 µm with a metal-OH band around 2.2-446
2.3 µm is a typical spectral feature of hydrated phyllosilicates (Fig. 4D). The best match here 447
is obtained for a Mg-rich trioctahedral smectite, such as saponite, which is consistent with the 448
composition of our initial olivine. The small absorption band at 2.39 µm visible in the 449
reference spectrum of saponite is also found in the spectra of the “Air” and “Air-H2O2” 450
samples, confirming the good match. The combination of the 1.9-µm and 2.31-µm bands is 451
also observed in the “CO2” sample, but this time the correlation is not as good as previously 452
in terms of intensity, because the 1.9-µm band is deeper than for the “Air” and “Air-H2O2” 453
samples, whereas the 2.31-µm one is shallower. Therefore, it is probable that this portion of 454
the spectrum of the “CO2” sample corresponds to the superimposition of the signature of a 455
Mg-rich smectite with the one of a distinct Mg-poor hydrated phase. 456
Another wavelength range of interest is the one located between 3.8 and 4.2 µm, 457
because this is where the main absorption bands related to carbonate minerals should appear 458
(Fig. 4E and F). However, such carbonate-related bands are not obvious in the spectra of our 459
weathered samples (Fig. 4C). Indeed, the small 4.02-µm band found in the spectrum of the 460
initial olivine is still present without significant change, except perhaps a subtle enlargement 461
toward the short wavelengths for the “CO2” sample. This indicates that no (or nearly no) 462
carbonates were formed during the experiment, whatever the weathering conditions. 463
NIR spectra were also obtained from the <2-µm fraction of the weathered powders and 464
are reported in Fig. 6. These spectra are in good agreement with those of the bulk samples 465
(Fig. 4): the spectral signature of olivine appears preserved, while absorption bands related to 466
20
secondary phases are obvious at 1.91 and 2.31 µm. Moreover, the latter is the shallowest for 467
the “CO2-H2O2” sample, where it is probably inherited from the initial olivine. The intensity 468
of the 2.31-µm band then increases for the “CO2”, “Air-H2O2” and “Air” samples, 469
respectively. This confirms the trend shown in Fig. 5. Finally, a small absorption band located 470
at 1.39 µm is visible in the four spectra, while it was absent from Fig. 4. However, its 471
presence is not surprising, since it is classically associated to the one at 1.91 µm in hydrated 472
samples (Hunt, 1977). 473
474
3.2.2. X-ray diffraction 475
476
No clay signal appeared in the diffraction patterns of the weathered bulk samples (data 477
not shown), nor in those of the air-dried oriented <2-µm fractions (Fig. 7). This suggests that 478
the smectite phase detected in three samples by NIR spectroscopy has a low abundance and/or 479
a low crystallinity, both possibilities being consistent with the timeframe of the experiment. 480
481
3.2.3. Transmission electron microscopy 482
483
Grains from the initial olivine observed by TEM show clean and smooth surfaces, 484
which are the result of grinding (Fig. 8A). In contrast, and although some unmodified surfaces 485
persisted, the formation of two new phases directly on the olivine grains is clearly observed in 486
the weathered samples. The first type is characterized by its “cotton-like” texture (Fig. 8B) 487
and the second type by its “filamentous” texture (Fig. 8C). 488
The newly-formed phase with a cotton-like texture is observed only in the two 489
samples weathered under CO2. At low magnification, this phase appears as a semi-transparent 490
layer that surrounds a whole olivine grain, or even a group of several grains. Its thickness is 491
21
usually comprised between 50 and 100 nm, but can reach 400 nm in some cases. Its edge is 492
often darker, i.e., more opaque. At higher magnification, it turns out to be composed of small 493
rounded shapes difficult to separate visually, which gives the so-called cotton-like texture. 494
The newly-formed phase with a filamentous texture is abundant in the two samples 495
weathered under terrestrial air, where it appears on most olivine grains. It is also observed in 496
the “CO2” sample – although less frequently (<10% of the grains) –, but not in “CO2-H2O2”. 497
At low magnification, this phase gives a “hairy” aspect to the olivine grains. Its thickness is 498
never above a few tens of nanometers. At higher magnification, it becomes possible to 499
distinguish individual filaments, which are pointing more or less outward, depending on the 500
observed grain. 501
The good correlation between the band depth at 2.31 µm in the NIR spectra and the 502
abundance of the filamentous phase in TEM imagery suggests that this phase corresponds to 503
the smectite phase. As highlighted in Fig. 8D, this is further confirmed by a comparison with 504
phyllosilicates described in the literature, which have very similar morphologies (e.g., Smith 505
et al., 1987; Giorgetti et al., 2001; Jones and Brearley, 2006; Bishop et al., 2007; Lantenois et 506
al., 2008). Unfortunately, it was not possible to directly measure the distance between 507
smectite layers by high-resolution TEM because this phase was highly unstable under the 508
electron beam. 509
In order to precisely determine the chemistry of the newly-formed phases, EDX 510
measurements were achieved on the “CO2” and “Air” samples. The initial olivine was also 511
analyzed to obtain a reference composition. The resulting dataset is presented here in the form 512
of a ternary diagram Si-Mg-Fe (Fig. 9). The mean compositions calculated for each phase are 513
available in the electronic annex (Table S2). 514
The compositions measured by EDX for grains of the initial olivine are 515
homogeneously distributed around the bulk composition determined by ICP-OES, proving a 516
22
good agreement between the two methods (Fig. 9A and B). In contrast, the data points 517
corresponding to the newly-formed phases are not concentrated around a precise composition, 518
but rather distributed along mixing lines between the initial olivine and the secondary phase 519
itself (Fig. 9C, D and E). For the filamentous phase, this chemical mixing is explained by the 520
irregularity and low thickness of the weathered layer and by its direct proximity with the 521
unaltered core of the olivine grain: in these conditions, it is difficult to retrieve signal from the 522
secondary phase only. For the cotton-like phase, which is easier to analyze due to its greater 523
thickness, the mixing line may either reflect different stages in the alteration process or be due 524
to the presence of remaining olivine fragments within the weathered layer. 525
The filamentous phase is chemically closer to the initial olivine compared to the 526
cotton-like phase. Indeed, although both are characterized by a loss of Mg relative to Si, this 527
loss is much more pronounced for the cotton-like phase. 528
Compositions measured for the filamentous phase are compatible with a Mg/Fe-529
bearing smectite, intermediate between the saponite and nontronite endmembers. However, 530
there are some differences between the filamentous phase of the “Air” sample and the one of 531
the “CO2” sample. In the first case, the loss of Fe is important and so, the data points tend 532
toward a Mg-rich smectite close to the saponite endmember (Fig. 9A and C). In the second 533
case, the loss of Fe is less pronounced and some points even indicate an enrichment in Fe 534
(Fig. 9A and D). Hence, the corresponding smectite would be more Fe-rich than in the “Air” 535
sample, although still closer to saponite than nontronite. These observations are confirmed by 536
the mean compositions (Table S2): the filamentous phase of “Air” sample lost 21% of Mg and 537
24% of Fe relative to the initial olivine, while its counterpart in the “CO2” sample lost 42% 538
and 7%, respectively. 539
The data points measured for the cotton-like phase clearly tend toward the Si apex of 540
the ternary diagram (Fig. 9A and E). This reflects a major loss of both Mg and Fe relative to 541
23
the initial olivine (76% and 60%, respectively; Table S2). The most likely single phase with 542
such Si-rich composition – Si/(Si+Mg) > 0.8 – is silica (SiO2·nH2O) because most 543
phyllosilicate minerals have Si/(Si+Mg) < 0.7. Furthermore, the cotton-like phase is 544
composed of rounded shapes similar to the spheres observed in natural opals, although the 545
latter are usually slightly bigger (~150-350 nm; e.g., Jones and Segnit, 1969; Rondeau et al., 546
2004). 547
Table S2 also shows the behavior of some minor elements. For example, Na, Cl and K 548
are always less abundant in the secondary phases compared to the initial olivine, which is 549
consistent with their high solubility. On the contrary, the Al content is always higher in the 550
secondary phases, which reflects its low solubility. 551
552
3.2.4. Mössbauer spectroscopy 553
554
The four types of atmospheres used in our experiment cover a large span of redox 555
states, from reducing conditions (CO2 atmosphere) to moderately oxidizing (terrestrial 556
atmosphere) to highly oxidizing (H2O2). Analyses of the solutions have shown that, as 557
expected, only the reducing conditions allowed ferrous iron to stay dissolved. In contrast, in 558
the three other conditions, any Fe2+ extracted from olivine must have been converted to Fe3+ 559
and incorporated to (oxy)hydroxides or other Fe3+-bearing phase. We acquired Mössbauer 560
spectra of our samples (Fig. S3; included in the electronic annex) in an attempt to identify and 561
quantify such ferric phases (e.g., Schröder et al., 2004; and references therein). However, in 562
all the samples, we obtained only the signature of olivine, through the absorptions of Fe2+. 563
The absence of Fe3+ detection gives an upper limit of the order of a few percent. This shows 564
that the addition of hydrogen peroxide has not increase the oxidation of Fe2+ in measurable 565
proportions compared to the other environments tested here. 566
24
567
3.2.5. Thermogravimetry and total inorganic carbon 568
569
A primary mineral like olivine contains negligible volatiles: as a consequence, no 570
mass loss is expected in a thermogravimetric analysis, and this is verified here (Fig. 10). In 571
contrast, during weathering, secondary phases incorporate volatiles, including water (in the 572
form of H2O molecules and/or hydroxyl groups OH) in the case of smectites or silica and CO2 573
in the case of carbonates. 574
Differential thermogravimetric (DTG) curves obtained for the “CO2-H2O2” and “CO2” 575
samples are similar (Fig. 10): they show mass losses centered at ~100 °C and ~410 °C. In 576
addition, the “CO2” sample presents two minor mass losses at ~550 °C and ~630 °C. In 577
comparison, the curve of the “Air” sample is closer to the one of the initial olivine: its shows a 578
unique mass loss centered at ~80 °C. Note here that the “Air-H2O2” sample has not been 579
analyzed due to its high similarity with the “Air” sample, as revealed by previous results. 580
Mass losses located at 100 °C or less are attributable to the release of weakly-bound 581
water (e.g., Ek et al., 2001). Having in mind the NIR and TEM results, this water probably 582
comes essentially from hydrated silica for the samples weathered under CO2 and from the 583
smectite phase for the “Air” sample. A small contribution from the smectite phase is also 584
expected for the “CO2” sample. 585
Mass losses located at 410 °C for the “CO2-H2O2” and “CO2” samples are attributable 586
to the release of hydroxyl groups. Several alteration phases could have produced them. For 587
example, brucite Mg(OH)2 undergoes dehydroxylation at this temperature (Kotra et al., 1982), 588
but this hypothesis can be ruled out because it is not consistent with the NIR spectra (Fig. 4 589
and 6): if present, brucite would produce a strong absorption band at 1.4 µm. The thermal 590
decomposition of Fe-hydroxides is more plausible. Indeed, although goethite and 591
25
lepidocrocite undergo dehydroxylation at less than 300 °C, it has been shown that the 592
decomposition temperature is above 400 °C for hydrated forms close to ferrihydrite (Prasad 593
and Sitakara Rao, 1984; Mitov et al., 2002). However, no Fe-(oxy)hydroxides has been 594
formally identified by NIR, XRD or TEM. A third possibility is the dehydroxylation of 595
hydrated silica, which occurs at higher temperature than its dehydration (Ek et al., 2001). This 596
last hypothesis would be consistent with the detection of an abundant Si-rich phase by TEM 597
and with the fact that the mass losses at 100 °C and 410 °C appear to be correlated in the 598
curves of the two samples containing this phase. 599
Finally, minor mass losses at 550 °C and 630 °C observed only for the “CO2” sample 600
can be attributed to the release of CO2, which implies the presence of minor amounts of 601
carbonates in this powder (e.g., Kotra et al., 1982). This detection is coherent with the very 602
small amount of TIC measured in this sample (~0.09 mg/g), in contrast with null values for 603
the initial olivine and the other weathered samples. Since two distinct gas releases are 604
observed, they must correspond to two distinct mineral phases, for example siderite and 605
magnesite, respectively (Kotra et al., 1982). Because the direct precipitation of magnesite in 606
our experimental conditions is unlikely (Hänchen et al., 2008), hydrated Mg-carbonates such 607
as nesquehonite or hydromagnesite may have formed instead and then been converted to 608
magnesite during the thermogravimetric analysis. 609
Combining the above qualitative interpretations with the quantitative mass losses 610
observed – in addition to results from other analytical methods – allows to estimate directly or 611
indirectly the abundances of the secondary phases, as summarized in Table 2. 612
613
4. DISCUSSION 614
615
4.1. Nature and abundance of secondary phases 616
26
617
4.1.1. Smectite phase 618
619
The formation of a Mg-rich clay in the samples weathered under terrestrial air and in 620
lower amount in the “CO2” sample is attested by NIR spectra, TEM imagery and EDX 621
analyses (Table 2). Although definitive identification of the exact species by XRD was not 622
possible, the NIR signature is best matched by a Mg-rich member of the smectite group, 623
probably close to saponite (Fig. 4 and 6). On the contrary, the NIR spectra of our samples lack 624
features typical of some other Mg-bearing clays, such as talc (sharp and deep band at 1.39 625
µm), montmorillonite (2.21- or 2.29-µm band), palygorskite (2.2-µm band) and serpentine 626
(2.1-µm band). Serpentine can be further discounted because it does not match the chemical 627
compositions retrieved by EDX (Fig. 9). 628
We proposed in section 3.2.2 that the non-detection of the smectite phase by XRD – 629
even using oriented <2-µm mounts – could be due a low abundance and/or a low crystallinity 630
of the smectite. Consistently, TEM observations have shown that the filamentous phase was 631
really thin (several tens of nanometers at most). In addition, filaments were never seen 632
isolated from the olivine grains; therefore, they may have been unable to deposit flat, which 633
would have in turn limited the intensity of the (001) reflections of clay particles in the XRD 634
analyses. Finally, the hypothesis of a low crystallinity remains plausible, given the limited 635
timeframe of the experiment. As a consequence, the filamentous phase may not be a mature 636
smectite, but rather an early stage of crystallization or “precursor” (e.g., Santiago Buey et al., 637
2000). 638
Finally, a possibly important result of our study is that the smectite phase has been 639
observed in the “CO2” sample but not in the “CO2-H2O2” sample. This suggests that hydrogen 640
27
peroxide had an inhibiting effect on this secondary phase. This point will be further discussed 641
in section 4.2.2. 642
643
4.1.2. Si-rich phase 644
645
TEM observations of the “CO2” and “CO2-H2O2” samples have revealed a newly-646
formed phase with a so-called cotton-like texture, which is absent from the powders 647
weathered under terrestrial atmosphere. EDX analyses have shown that this phase clearly 648
tends toward a composition of silica (SiO2·nH2O). The formation of such Si-rich phase is 649
consistent with the observed evolution of Si concentrations in solution. Indeed, under CO2, Si 650
concentrations stabilized after 14 or 31 days, what may reflect equilibrium with a solid phase 651
(Fig. 3). Because of the low abundance of the smectite phase in “CO2” and its absence in 652
“CO2-H2O2”, it cannot explain the intensity of the hydration band at 1.9 µm in the NIR 653
spectra of these two samples (Fig. 5), nor the H2O-related mass losses in their 654
thermogravimetric analyses (Fig. 10). Therefore, hydration is most probably borne by the Si-655
rich phase, consistently with the formula of SiO2·nH2O. 656
From the Si-Mg-Fe ternary diagram (Fig. 9), the filamentous (smectite) and cotton-657
like (Si-rich) phases of the “CO2” sample seem to follow a same trend between the initial 658
olivine and the Si apex. This could suggest that they represent two different stages of a unique 659
alteration process of olivine. Interestingly, King et al. (2010) have reported the transformation 660
of a Si-rich phase into lizardite during the alteration of olivine under CO2 at 200 °C. A similar 661
mechanism may have occurred in our experiment, but the trend in Fig. 9 may also be 662
consistent with the reverse, i.e. the transformation of the smectite phase into the Si-rich phase. 663
In this case, the coexistence of the two phases would be the result of differential weathering 664
(in intensity or in time) within the powder. However, it should be noted that no TEM 665
28
observation (for example, of a potential intermediate morphology) supports either of these 666
hypotheses. Thus, alternatively, the smectite phase and the Si-rich phase could represent two 667
different reaction pathways from the initial olivine. In all cases, it is clear that the powder 668
grains have not all undergone the same alteration, or at least not with the same intensity. This 669
is further confirmed by the observation by TEM of grains with clean surfaces, obviously 670
intact despite the three months in the water bath (Fig. 8). This differential weathering suggests 671
that very low-scale (micrometers, or even nanometers) processes may have a strong influence, 672
and that interfacial conditions (pH, dissolved elements, L/R ratio) are not necessarily 673
representative of the bulk solution (e.g., King et al., 2010; Daval et al., 2011). 674
675
4.1.3. Fe-(oxy)hydroxides 676
677
No Fe-(oxy)hydroxide has been formally identified in our weathered samples by any 678
of the analytical methods employed and no reddening of the weathered powders has been 679
observed. This implies that such phase, if formed, must be limited to very minor amounts. 680
However, EDX data of the secondary phases (Fig. 9 and Table S2) provides some information 681
above the behavior of Fe during experimental weathering. For example, the Si-rich phase 682
formed in the “Mars reactors” has lost a significant amount of Fe compared to the olivine 683
from which it derives (Table S2). This means that the “missing” Fe must have been dissolved 684
or incorporated in a Fe-rich phase. In the “CO2-H2O2” reactor, Fe2+ released from olivine is 685
expected to be immediately oxidized and precipitated as Fe-(oxy)hydroxides, which is 686
consistent with its non-detection in the solution (Fig. 3). In the “CO2” reactor, reducing 687
conditions has allowed Fe2+ to stay in solution. However, the progressive decrease of Fe 688
concentration after 14 days (Fig. 3) along with the detection of H2 in the headspace suggest 689
that partial oxidation occurred through the dissociation of H2O molecules, possibly forming 690
29
some minor Fe-(oxy)hydroxides (Table 2). Alternatively, Fe may have been incorporated in 691
other Fe2+ or Fe3+-bearing phases, such as the smectite phase or siderite. 692
693
4.1.4. Carbonates 694
695
No carbonate has been formally identified in our samples by NIR spectroscopy, 696
despite the method being very sensitive to this group of minerals (e.g., Gaffey, 1986; Bishop 697
et al., 2001). Only a subtle enlargement toward the short wavelengths of the pre-existing 4.02-698
µm band was observed in the spectrum of the “CO2” sample (Fig. 4), suggesting the very 699
minor appearance of a carbonate mineral. Then, thermogravimetric analyses have revealed 700
traces of CO2 trapped in the same sample, which was confirmed by the TIC measurement. 701
The formation of siderite is possible, given the absence of O2 in this reactor. In any case, the 702
abundance involved must be very low (Table 2). 703
704
4.2. Effects of environmental parameters and implications for early Mars 705
706
Due to unavoidable differences in extrinsic (scale, timeframe) and intrinsic factors 707
(grain size, composition), laboratory experiments never exactly mimic natural processes (e.g., 708
Casey et al., 1993). For example, pedogenic processes – i.e., vertical mobility of dissolved 709
elements and pH variations within weathering profiles – are not reproduced in our 710
experimental setting, although they have a primary influence on the nature and abundance of 711
secondary phases (e.g., Gaudin et al., 2011). Nevertheless, our results allow us to isolate the 712
effects of several key environmental parameters on weathering processes, which are reviewed 713
below. 714
715
30
4.2.1. Effects of CO2 716
717
Our results emphasize the important effects of a CO2 atmosphere and associated low 718
pH on the weathering products of forsteritic olivine. Under terrestrial atmosphere, the main 719
(and only) secondary phase found is a Mg-rich smectite. In contrast, under CO2, the main 720
secondary phase found is a hydrated, Si-rich phase. However, an important result of the study 721
is that a CO2 atmosphere does not prevent the formation of a Mg-rich smectite from olivine. 722
The reasons for which the smectite phase is in lower abundance under CO2 compared to 723
terrestrial conditions (Table 2) are discussed below and in the next section. 724
The development of a Si-rich layer has been reported in numerous experiments of 725
olivine carbonation, even when the bulk solution was undersaturated with respect to 726
amorphous silica (e.g., Pokrovsky and Schott, 2000; Bearat et al., 2006; Garcia et al., 2010; 727
King et al., 2010; Daval et al., 2011). Such Si-rich phase is often described as a “passivating 728
layer”, which tends to decrease the dissolution rate of olivine and, in some cases, prevents the 729
formation of any other secondary phase (Giammar et al., 2005; Daval et al., 2011). This effect 730
would have been especially effective here, because the discontinuous, manual stirring of our 731
reactors is not expected to have been vigorous enough to remove such a passivating layer 732
from the grains during the course of the experiment. Therefore, it is possible that the 733
development of the Si-rich phase in the “Mars samples” have contributed – in addition to the 734
role of Fe discussed in the next section – to decrease the abundance of the smectite phase 735
compared to the “Earth samples”. 736
This same passivating effect of the Si-rich phase may also have limited the amount of 737
carbonates formed in our experiment. Indeed, the very low amount of carbonates is a major 738
discrepancy with other studies ran at high pressures of CO2 and/or higher temperatures (e.g., 739
Pokrovsky and Schott, 2000; Gerdemann et al., 2007; Garcia et al., 2010; King et al., 2010), 740
31
with the notable exception of the work of Giammar et al. (2005) and Daval et al. (2011) 741
mentioned previously. Interestingly, this low abundance of carbonates is consistent with the 742
overall secondary mineralogy of the martian surface, which is dominated by Fe/Mg smectites 743
and shows only rare and local occurrences of carbonates (Ehlmann et al., 2008b; Carter et al., 744
2013). 745
Taken together, the present results and previous studies of olivine carbonation are 746
consistent with a high production of silica through weathering of olivine-bearing primary 747
rocks on early Mars. This could possibly explain the high-silica rock compositions measured 748
within the northern plains (e.g., McLennan, 2003), as well as the recent detection of opaline 749
silica in numerous alluvial and delta fans (Carter et al., 2012). However, it is worth noting that 750
the Si-rich phase formed in our experiment has no diagnostic signature in the NIR domain 751
(Fig. 4), which implies that similar materials could be difficult to identify by orbital 752
spectroscopy. 753
Finally, a dense CO2 atmosphere implies more acidic meteoric water than on present-754
day Earth. As highlighted in section 3.1 and previous studies (e.g., Olsen and Rimstidt, 2007; 755
and references therein), this greatly enhances the dissolution rate of olivine. As a 756
consequence, a CO2 atmosphere would favor the leaching of soluble elements – Mg in 757
particular, which is ~30 times more abundant in the “CO2” solution than in the “Air” solution 758
at the end of the experiment (Fig. 3). Such an intense leaching would accelerate the 759
development of weathering profiles and facilitate the appearance of an upper Al-rich, 760
kaolinite-bearing layer from an Al-poor bedrock, without requiring (highly unlikely) tropical 761
conditions as on Earth (Gaudin et al., 2011). Moreover, a CO2 atmosphere – with low 762
production or shallow extinction depth of H2O2 (Zent, 1998) – would maintain reducing 763
conditions within these weathering profiles and thus increase the mobility of Fe (e.g., 764
Murakami et al., 2004). In such conditions, the upper zone of putative martian weathering 765
32
profiles would be less rich in Fe than their terrestrial counterparts (Gaudin et al., 2011; 766
Greenberger et al., 2012) and Al would be the only major element to remain “immobile”, 767
along with some Si to produce kaolinite. Therefore, the low Fe-oxide content generally 768
observed by remote instruments in Al-rich units on Mars does not ruled out pedogenesis as a 769
formation process for Al-rich clays over Fe/Mg-smectites superimpositions. 770
771
4.2.2. Effects of H2O2 772
773
Hydrogen peroxide was used in this experiment to represent the highly oxidizing 774
compounds formed by photochemical reactions in the Mars atmosphere (Clancy et al., 2004; 775
Encrenaz et al., 2004) and detected in the regolith by the Viking landers (Bullock et al., 1994; 776
Zent, 1998), which may have been more abundant on a “wetter” early Mars. Its effect on the 777
dissolution of olivine turns out to be weak in the timeframe of the experiment, as highlighted 778
by the ICP-OES analyses (Fig. 3). Similarly, Mössbauer spectra (Fig. S3) indicate that H2O2 779
has not enhanced the oxidation of our Mg-rich olivine in a measurable way. 780
On the other hand, we have noticed an important effect of this molecule regarding the 781
smectite phase, which formed in the “CO2” reactor but not in the “CO2-H2O2” one. Thus, 782
H2O2 seems to inhibit the formation of smectite under CO2 conditions. The only difference 783
between the two “Mars reactors” is the redox state of the system. Therefore, a hypothesis to 784
consider is that the inhibiting effect is related to the rapid oxidation of Fe2+ deriving from the 785
dissolution of olivine (as shown by Fig. 3, Fe2+ is present in the solution of the “CO2” reactor 786
but not in the one of the “CO2-H2O2” reactor, which suggests the precipitation of Fe-787
(oxy)hydroxides in the latter). Previous studies have shown that Fe2+ plays a strong role in the 788
formation of smectite at the low pH values associated with CO2 conditions (Murakami et al., 789
2004; Tosca et al., 2008). At pH ~4.6, Murakami et al. (2004) observed that Fe2+-rich 790
33
vermiculite or smectite precipitated at the edge of their Fe-bearing biotite under anoxic 791
conditions, while only Fe3+- and Al-(hydr)oxides were observed under oxic conditions. 792
Consistently, Tosca et al. (2008) reported that in anoxic conditions, Fe2+-bearing saponite 793
precursor can precipitate at pH as low as 5, whereas pure Mg-saponite form only at pH 9. 794
Moreover, this result is consistent with NIR spectra obtained by Dehouck et al. (2012) in their 795
weathering experiment conducted at low temperature and low L/R ratio. Indeed, two of their 796
silicate samples (olivine Ol1 and orthopyroxene OPx) weathered under CO2 showed a small 797
Mg-OH absorption band at 2.31 µm – attributable to a Mg-rich smectite phase, although not 798
confirmed due to the lack of TEM data – while the same samples weathered under CO2+H2O2 799
did not. 800
These observations imply that Fe2+ is required to ensure the stability of the smectite 801
phase under CO2 conditions (pH ~6 in our experiment) and suggest that H2O2 has prevented 802
its incorporation into this phase in the case of the “CO2-H2O2” reactor. Since our starting 803
olivine was relatively Fe-poor, this also explains the low abundance of the smectite phase in 804
the “CO2” sample (compared to “Air” and “Air-H2O2”), although the Si-rich phase may also 805
play a role here as discussed previously. This is also consistent with the EDX analyses, which 806
show that the filamentous phase of the “CO2” sample is richer in Fe than its counterpart 807
formed in the “Earth reactors” (Fig. 9). Hence, CO2 conditions seem to favor the formation of 808
Fe-bearing smectites rather than purely Mg-bearing ones. Therefore, Fe-rich starting material 809
(in laboratory experiments) or bedrock (in natural weathering) would probably tend to 810
produce proportionally more smectites than in this study. 811
Under terrestrial air, the role of Fe is less important than under CO2 because the pH is 812
higher (Fig. 3). Thus Fe is not required to ensure the stability of the smectite phase: as a 813
result, the smectite phase is more abundant and the effect of hydrogen peroxide is less 814
pronounced. However, a small inhibiting effect of H2O2 in terrestrial conditions cannot be 815
34
completely discounted since the 2.31-µm absorption band attributed to the smectite phase is 816
slightly shallower for the “Air-H2O2” sample than for the “Air” sample (Fig. 5 and 6). 817
Taken together, these results suggest that if the highly oxidizing compounds found on 818
present-day Mars were more abundant 3 to 4 billion years ago, they would have been an 819
obstacle for the formation of abundant clay minerals. However, images of the first drilling 820
operations performed by the Curiosity rover at Gale crater have shown that the smectite-821
bearing sediments of Yellowknife Bay are gray-colored at <2 cm depth (as shown by the 822
“mini-drill” of sol 180), contrasting with the reddish surface and suggesting only very shallow 823
oxidation at this place (Grotzinger et al., 2014; Vaniman et al., 2014). In addition, the 824
detection of reduced sulfur in these sediments is another indication of limited oxidation 825
(Grotzinger et al., 2014). The presence of abundant clays implies either that the production of 826
highly oxidizing compounds was limited on early Mars or that their inhibiting effect on the 827
formation of clays was counterbalanced by other mechanisms (e.g., Fe-rich source rock and/or 828
higher pH; Vaniman et al., 2014). Nevertheless, the role of H2O2 may have increase during 829
the subsequent eras (Dehouck et al., 2012) and ultimately become dominant (Huguenin, 1982; 830
Bibring et al., 2006) due to the progressive disappearance of other forms of weathering (i.e., 831
those involving liquid water). In addition, long term exposure of initial Fe2+-bearing smectites 832
to H2O2 may have cause their gradual transformation into Fe3+-bearing smectites (Beehr and 833
Catalano, 2012). 834
835
4.2.3. Effects of L/R ratio 836
837
In a previous weathering experiment, Dehouck et al. (2012) used several silicate 838
samples, including olivine with a composition similar to the one used here (Fo90Fa10). Thus, it 839
is possible to directly compare the results of these two experiments, which differ mainly by 840
35
their L/R ratio, low in Dehouck et al. (2012) (thin films of water condensed on the olivine 841
grains) while relatively high in the present study (30 g of olivine powder in 300 mL of water). 842
This gives the opportunity to evaluate the effect of the L/R ratio on the weathering of olivine 843
under CO2 conditions. 844
The only secondary phase formed from olivine formally identified in Dehouck et al. 845
(2012) was nesquehonite, a hydrated Mg-carbonate, along with possible Mg-smectite phase in 846
the samples weathered without H2O2, as mentioned above. In contrast, in the present study, 847
silica (along with smectite phase in the case of the “CO2” sample) dominates over carbonates. 848
Therefore, it appears that the L/R ratio has a major effect on the nature of the secondary 849
phases formed from olivine under CO2, with low ratios favoring carbonates and higher ratios 850
favoring silica. This is broadly consistent with results obtained at 150 °C and 150 bar by 851
Garcia et al. (2010), which obtained more carbonates for a L/R ratio of 0.1 than for a ratio of 852
10. Moreover, Bearat et al. (2006) noted that, for stirred experiments, lower L/R ratios can 853
lead to higher particle-particle abrasion, which in turn contribute to remove the Si-rich layer 854
and thus increase the formation of carbonates. 855
856
5. CONCLUSION 857
858
The laboratory experiment described in this paper has allowed us to compare the 859
weathering of olivine in simulated early martian conditions and terrestrial conditions, and to 860
evaluate the effect of hydrogen peroxide on this process. Our results show that the type of 861
atmosphere clearly has a primary influence on weathering pathways. The main (and only) 862
secondary phase formed under terrestrial air is a Mg-rich smectite phase. Under CO2, a 863
smectite phase appeared as well, but only in the absence of H2O2. Therefore, a pure CO2 864
atmosphere does not prevent the formation of smectites from olivine, although the addition of 865
36
highly oxidizing compounds could inhibit the process. However, the main secondary phase 866
formed under CO2 turns out to be a silica-rich phase with a non-diagnostic near-infrared 867
spectral signature. Finally, carbonate minerals are absent or nearly absent in all of our final 868
samples. 869
Several regions of Mars, such as Nili Fossae, are known to host olivine-rich bedrock 870
(Hoefen et al., 2003; Mustard et al., 2005; Poulet et al., 2007; Ody et al., 2013) and to have 871
undergone rock-water-atmosphere interactions (Ehlmann et al., 2008a; Gaudin et al., 2011). 872
Fe/Mg-smectites are the most common secondary phases found in these regions (e.g., 873
Mangold et al., 2007; Ehlmann et al., 2011). In Gale crater, the Curiosity rover has recently 874
found evidence for low grade diagenesis of lacustrine sediments, which resulted in the 875
formation of Mg-bearing smectite clays at the expense of olivine (Grotzinger et al., 2014; 876
McLennan et al., 2014; Vaniman et al., 2014). In these two contexts, our experimental results 877
are helpful to understand the weathering processes that occurred. They show that Fe/Mg-878
smectite clays can be formed at the low pH associated with a dense CO2 atmosphere and 879
highlight the critical importance of the redox conditions, the Fe-content of the source rock and 880
the formation of a passivating layer around the altered grains. 881
882
Acknowledgments. The authors thank two anonymous reviewers and the associate 883
editor Penelope King for their comments and suggestions which greatly improved the quality 884
of the manuscript. The authors alsothank all the people who brought them their scientific or 885
technical expertise and without whom this study would not have been possible: Hervé Loyen, 886
Laurent Lenta and Carole La from the LPGN laboratory (Nantes, France); Pierre-Emmanuel 887
Petit, Philippe Leone, Michel Suchaud, Nicolas Stéphant and Stéphane Grolleau from the 888
IMN laboratory (Nantes, France); Guillaume Blain, Francis Crumière and Massoud Fattahi-889
Vanani from the Subatech laboratory (École des Mines, Nantes, France); Damien Chaudanson 890
37
from the CINaM laboratory (Marseille, France); and Patrick Le Bescop from the CEA 891
laboratory (Saclay, France). Thanks to Steven Jaret for proofreading the manuscript. This 892
work was supported by the Centre National de la Recherche Scientifique (CNRS) through its 893
EPOV (Environnements Planétaires et Origines de la Vie) and PNP (Programme National de 894
Planétologie) programs, and by the Centre National d’Études Spatiales (CNES). 895
896
REFERENCES 897
898
Amram, K., Ganor, J., 2005. The combined effect of pH and temperature on smectite 899
dissolution rate under acidic conditions. Geochimica et Cosmochimica Acta 69, 2535-900
2546. 901
Ansan, V., Loizeau, D., Mangold, N., Le Mouélic, S., Carter, J., Poulet, F., Dromart, G., 902
Lucas, A., Bibring, J. P., Gendrin, A., Gondet, B., Langevin, Y., Masson, P., Murchie, 903
S., Mustard, J. F., and Neukum, G., 2011. Stratigraphy, mineralogy, and origin of 904
layered deposits inside Terby crater, Mars. Icarus 211, 273-304. 905
Bauer, A., Schafer, T., Dohrmann, R., Hoffmann, H., Kim, J.I., 2001. Smectite stability in 906
acid salt solutions and the fate of Eu, Th and U in solution. Clay Minerals 36, 93-103. 907
Bearat, H., McKelvy, M. J., Chizmeshya, A. V. G., Gormley, D., Nunez, R., Carpenter, R. 908
W., Squires, K., and Wolf, G. H., 2006. Carbon sequestration via aqueous olivine 909
mineral carbonation: Role of passivating layer formation. Environmental Science & 910
Technology 40, 4802-4808. 911
Beehr, A. R. and Catalano, J. G., 2012. Oxidation pathways of ferrous iron phyllosilicates: 912
insights into Martian phyllosilicate formation. Third Conference on Early Mars: 913
Geologic, Hydrologic and Climatic Evolution, abstract #7008. 914
38
Bibring, J. P., Langevin, Y., Gendrin, A., Gondet, B., Poulet, F., Berthe, M., Soufflot, A., 915
Arvidson, R., Mangold, N., Mustard, J., and Drossart, P., 2005. Mars surface diversity 916
as revealed by the OMEGA/Mars Express observations. Science 307, 1576-1581. 917
Bibring, J. P., Langevin, Y., Mustard, J. F., Poulet, F., Arvidson, R., Gendrin, A., Gondet, B., 918
Mangold, N., Pinet, P., and Forget, F., 2006. Global mineralogical and aqueous Mars 919
history derived from OMEGA/Mars express data. Science 312, 400-404. 920
Bishop, J. L., Lougear, A., Newton, J., Doran, P. T., Froeschl, H., Trautwein, A. X., Korner, 921
W., and Koeberl, C., 2001. Mineralogical and geochemical analyses of Antarctic lake 922
sediments: A study of reflectance and Mossbauer spectroscopy and C, N, and S 923
isotopes with applications for remote sensing on Mars. Geochimica et Cosmochimica 924
Acta 65, 2875-2897. 925
Bishop, J. L., Schiffman, P., Murad, E., Dyar, M. D., Drief, A., and Lane, M. D., 2007. 926
Characterization of alteration products in tephra from Haleakala, Maui: A visible-927
infrared spectroscopy, Mossbauer spectroscopy, XRD, EMPA and TEM study. Clays 928
and Clay Minerals 55, 1-17. 929
Bishop, J. L., Tirsch, D., Tornabene, L. L., Jaumann, R., McEwen, A. S., McGuire, P. C., 930
Ody, A., Poulet, F., Clark, R. N., Parente, M., McKeown, N. K., Mustard, J. F., 931
Murchie, S. L., Voigt, J., Aydin, Z., Bamberg, M., Petau, A., Michael, G., Seelos, F. 932
P., Hash, C. D., Swayze, G. A., and Neukum, G., 2013. Mineralogy and morphology 933
of geologic units at Libya Montes, Mars: Ancient aqueously derived outcrops, mafic 934
flows, fluvial features, and impacts. Journal of Geophysical Research: Planets 118, 935
487-513. 936
Bullock, M. A. and Moore, J. M., 2007. Atmospheric conditions on early Mars and the 937
missing layered carbonates. Geophysical Research Letters 34. 938
39
Bullock, M. A., Stoker, C. R., McKay, C. P., and Zent, A. P., 1994. A coupled soil-939
atmosphere model of H2O2 on Mars. Icarus 107, 142-154. 940
Carter, J., Poulet, F., Bibring, J. P., Mangold, N., and Murchie, S., 2013. Hydrous minerals on 941
Mars as seen by the CRISM and OMEGA imaging spectrometers: Updated global 942
view. Journal of Geophysical Research: Planets, 1-29. 943
Carter, J., Poulet, F., Mangold, N., Ansan, V., Dehouck, E., Bibring, J.-P., and Murchie, S., 944
2012. Composition of alluvial fans and deltas on Mars. Lunar and Planetary Science 945
Conference XLIII, abstract #1978. 946
Casey, W. H., Banfield, J. F., Westrich, H. R., and McLaughlin, L., 1993. What do 947
dissolution experiments tell us about natural weathering? Chemical Geology 105, 1-948
15. 949
Chevrier, V., Mathe, P. E., Rochette, P., Grauby, O., Bourrie, G., and Trolard, F., 2006. Iron 950
weathering products in a CO2+(H2O or H2O2) atmosphere: Implications for weathering 951
processes on the surface of Mars. Geochimica et Cosmochimica Acta 70, 4295-4317. 952
Chevrier, V. F., Dehouck, E., Lozano, C. G., Altheide, T. S., 2012. Mineral parageneses 953
resulting from weathering on Early Mars and the effect of CO2 vs SO2 atmospheres. 954
Third Conference on Early Mars: Geologic, Hydrologic and Climatic Evolution, 955
abstract #7080. 956
Clancy, R. T., Sandor, B. J., and Moriarty-Schieven, G. H., 2004. A measurement of the 362 957
GHz absorption line of Mars atmospheric H2O2. Icarus 168, 116-121. 958
Clark, R. N., King, T. V. V., Klejwa, M., Swayze, G. A., and Vergo, N., 1990. High spectral 959
resolution reflectance spectroscopy of minerals. Journal of Geophysical Research-960
Solid Earth and Planets 95, 12653-12680. 961
40
Clark, R. N., G. A. Swayze, R. Wise, E. Livo, T. Hoefen, R. Kokaly, and S. J. Sutley (2007), 962
USGS digital spectral library splib06a, U.S. Geological Survey Digital Data Series 963
231. Available from: http://speclab.cr.usgs.gov/spectral.lib06 964
Daval, D., Sissmann, O., Menguy, N., Saldi, G. D., Guyot, F., Martinez, I., Corvisier, J., 965
Garcia, B., Machouk, I., Knauss, K. G., and Hellmann, R., 2011. Influence of 966
amorphous silica layer formation on the dissolution rate of olivine at 90 degrees C and 967
elevated pCO2. Chemical Geology 284, 193-209. 968
Dehouck, E., Chevrier, V., Gaudin, A., Mangold, N., Mathe, P. E., and Rochette, P., 2012. 969
Evaluating the role of sulfide-weathering in the formation of sulfates or carbonates on 970
Mars. Geochimica et Cosmochimica Acta 90, 47-63. 971
Dehouck, E., Mangold, N., Le Mouélic, S., Ansan, V., and Poulet, F., 2010. Ismenius Cavus, 972
Mars: A deep paleolake with phyllosilicate deposits. Planetary and Space Science 58, 973
941-946. 974
Delvigne, J., Bisdom, E. B. A., Sleeman, J., and Stoops, G., 1979. Olivines, their 975
pseudomorphs and secondary products. Pédologie 3, 247-309. 976
Ehlmann, B. L., Mustard, J. F., Murchie, S. L., Bibring, J. P., Meunier, A., Fraeman, A. A., 977
and Langevin, Y., 2011. Subsurface water and clay mineral formation during the early 978
history of Mars. Nature 479, 53-60. 979
Ehlmann, B. L., Mustard, J. F., Fassett, C. I., Schon, S. C., Head, J. W., Marais, D. J. D., 980
Grant, J. A., and Murchie, S. L., 2008a. Clay minerals in delta deposits and organic 981
preservation potential on Mars. Nature Geoscience 1, 355-358. 982
Ehlmann, B. L., Mustard, J. F., Murchie, S. L., Poulet, F., Bishop, J. L., Brown, A. J., Calvin, 983
W. M., Clark, R. N., Des Marais, D. J., Milliken, R. E., Roach, L. H., Roush, T. L., 984
Swayze, G. A., and Wray, J. J., 2008b. Orbital Identification of Carbonate-Bearing 985
Rocks on Mars. Science 322, 1828-1832. 986
41
Ehlmann, B. L., Mustard, J. F., Swayze, G. A., Clark, R. N., Bishop, J. L., Poulet, F., Marais, 987
D. J. D., Roach, L. H., Milliken, R. E., Wray, J. J., Barnouin-Jha, O., and Murchie, S. 988
L., 2009. Identification of hydrated silicate minerals on Mars using MRO-CRISM: 989
Geologic context near Nili Fossae and implications for aqueous alteration. Journal of 990
Geophysical Research-Planets 114. 991
Ek, S., Root, A., Peussa, M., and Niinisto, L., 2001. Determination of the hydroxyl group 992
content in Silica by thermogravimetry and a comparison with 1H MAS NMR results. 993
Thermochimica Acta 379, 201-212. 994
Encrenaz, T., Bézard, B., Greathouse, T. K., Richter, M. J., Lacy, J. H., Atreya, S. K., Wong, 995
A. S., Lebonnois, S., Lefèvre, F., and Forget, F., 2004. Hydrogen peroxide on Mars: 996
evidence for spatial and seasonal variations. Icarus 170, 424-429. 997
Fernandez-Remolar, D. C., Sanchez-Roman, M., Hill, A. C., Gomez-Ortiz, D., Prieto 998
Ballesteros, O., Romanek, C. S., and Amils, R., 2011. The environment of early Mars 999
and the missing carbonates. Meteoritics & Planetary Science 46, 1447-1469. 1000
Forget, F., Wordsworth, R., Millour, E., Madeleine, J. B., Kerber, L., Leconte, J., Marcq, E., 1001
and Haberle, R. M., 2012. 3D modelling of the early martian climate under a denser 1002
CO2 atmosphere: Temperatures and CO2 ice clouds. Icarus 222, 81-99. 1003
Gaffey, S. J., 1986. Spectral reflectance of carbonate minerals in the visible and near-infrared 1004
(0.35-2.55 microns): calcite, aragonite, and dolomite. American Mineralogist 71, 151-1005
162. 1006
Galan, E., Carretero, M. I., Fernandez-Caliani, J. C., 1999. Effects of acid mine drainage on 1007
clay minerals suspended in the Tinto River (Rio Tinto, Spain). An experimental 1008
approach. Clay Minerals 34, 99-108. 1009
Garcia, B., Beaumont, V., Perfetti, E., Rouchon, V., Blanchet, D., Oger, P., Dromart, G., Huc, 1010
A. Y., and Haeseler, F., 2010. Experiments and geochemical modelling of CO2 1011
42
sequestration by olivine: Potential, quantification. Applied Geochemistry 25, 1383-1012
1396. 1013
Gaudin, A., Dehouck, E., and Mangold, N., 2011. Evidence for weathering on early Mars 1014
from a comparison with terrestrial weathering profiles. Icarus 216, 257-268. 1015
Gerdemann, S. J., O'Connor, W. K., Dahlin, D. C., Penner, L. R., and Rush, H., 2007. Ex situ 1016
aqueous mineral carbonation. Environmental Science & Technology 41, 2587-2593. 1017
Giammar, D. E., Bruant, R. G., and Peters, C. A., 2005. Forsterite dissolution and magnesite 1018
precipitation at conditions relevant for deep saline aquifer storage and sequestration of 1019
carbon dioxide. Chemical Geology 217, 257-276. 1020
Giorgetti, G., Marescotti, P., Cabella, R., and Lucchetti, G., 2001. Clay mineral mixtures as 1021
alteration products in pillow basalts from the eastern flank of Juan de Fuca Ridge: a 1022
TEM-AEM study. Clay Minerals 36, 75-91. 1023
Golubev, S. V., Pokrovsky, O. S., and Schott, J., 2005. Experimental determination of the 1024
effect of dissolved CO2 on the dissolution kinetics of Mg and Ca silicates at 25 1025
degrees C. Chemical Geology 217, 227-238. 1026
Golubev, S. V., Bauer, A., Pokrovsky, O. S., 2006. Effect of pH and organic ligands on the 1027
kinetics of smectite dissolution at 25 °C. Geochimica et Cosmochimica Acta 70, 4436-1028
4451 1029
Greenberger, R. N., Mustard, J. F., Kumar, P. S., Dyar, M. D., Breves, E. A., and Sklute, E. 1030
C., 2012. Low temperature aqueous alteration of basalt: Mineral assemblages of 1031
Deccan basalts and implications for Mars. Journal of Geophysical Research 117, 1032
E00J12. 1033
Grotzinger, J., Sumner, D. Y., Kah, L. C., Stack, K., Gupta, S., Edgar, L., Rubin, D., Lewis, 1034
K., Schieber, J., Mangold, N., Milliken, R., Conrad, P. G., DesMarais, D., Farmer, J., 1035
Siebach, K., Calef III, F., Hurowitz, J., McLennan, S. M., Ming, D., Vaniman, D., 1036
43
Crisp, J., Vasavada, A., Edgett, K. S., Malin, M., Blake, D., Gellert, R., Mahaffy, P., 1037
Wiens, R. C., Maurice, S., Grant, J. A., Wilson, S., Anderson, R. C., Beegle, L., 1038
Arvidson, R., Hallet, B., Sletten, R. S., Rice, M., Bell III, J., Griffes, J., Ehlmann, B., 1039
Anderson, R. B., Bristow, T. F., Dietrich, W. E., Dromart, G., Eigenbrode, J., 1040
Fraeman, A., Hardgrove, C., Herkenhoff, K., Jandura, L., Kocurek, G., Lee, S., 1041
Leshin, L. A., Leveille, R., Limonadi, D., Maki, J., McCloskey, S., Meyer, M., Minitti, 1042
M., Newsom, H., Oehler, D., Okon, A., Palucis, M., Parker, T., Rowland, S., Schmidt, 1043
M., Squyres, S., Steele, A., Stolper, E., Summons, R., Treiman, A., Williams, R., 1044
Yingst, A. and the MSL Science Team, 2014. A habitable fluvio-lacustrine 1045
environment at Yellowknife Bay, Gale crater, Mars. Science 343, doi: 1046
10.1126/science.1242777. 1047
Hänchen, M., Prigiobbe, V., Baciocchi, R., and Mazzotti, M., 2008. Precipitation in the Mg-1048
carbonate system - effects of temperature and CO2 pressure. Chemical Engineering 1049
Science 63, 1012-1028. 1050
Hänchen, M., Prigiobbe, V., Storti, G., Seward, T. M., and Mazzotti, M., 2006. Dissolution 1051
kinetics of fosteritic olivine at 90-150 degrees C including effects of the presence of 1052
CO2. Geochimica et Cosmochimica Acta 70, 4403-4416. 1053
Hoefen, T. M., Clark, R. N., Bandfield, J. L., Smith, M. D., Pearl, J. C., and Christensen, P. 1054
R., 2003. Discovery of olivine in the Nili Fossae region of Mars. Science 302, 627-1055
630. 1056
Huguenin, R. L., 1982. Chemical weathering and the Viking biology experiments on Mars. 1057
Journal of Geophysical Research 87, 10069-10082. 1058
Hunt, G. R., Salisbury, J. W., and Lenhoff, C. J., 1972. Visible and near-infrared spectra of 1059
minerals and rocks: V. Halides, phosphates, arsenates, vanadates and borates. Modern 1060
Geology 3, 121-132. 1061
44
Hunt, G. R., 1977. Spectral signatures of particulate minerals in visible and near infrared. 1062
Geophysics 42, 501-513. 1063
Hurowitz, J. A., Fischer, W., Tosca, N. J., and Milliken, R. E., 2010. Origin of acidic surface 1064
waters and the evolution of atmospheric chemistry on early Mars. Nature Geoscience 1065
3, 323-326. 1066
Hurowitz, J. A., Tosca, N. J., McLennan, S. M., and Schoonen, M. A. A., 2007. Production of 1067
hydrogen peroxide in Martian and lunar soils. Earth and Planetary Science Letters 1068
255, 41-52. 1069
Jones, C. L. and Brearley, A. J., 2006. Experimental aqueous alteration of the Allende 1070
meteorite under oxidizing conditions: Constraints on asteroidal alteration. Geochimica 1071
et Cosmochimica Acta 70, 1040-1058. 1072
Jones, J. B. and Segnit, E. R., 1969. Water in sphere-type opal. Mineralogical Magazine 37, 1073
357-361. 1074
King, H. E., Plumper, O., and Putnis, A., 2010. Effect of Secondary Phase Formation on the 1075
Carbonation of Olivine. Environmental Science & Technology 44, 6503-6509. 1076
Knauth, L. P. and Epstein, S., 1982. The nature of water in hydrous silica. American 1077
Mineralogist 67, 510-520. 1078
Kotra, R. K., Gibson, E. K., and Urbancic, M. A., 1982. Release of volatiles from possible 1079
Martian analogs. Icarus 51, 593-605. 1080
Komadel, P., and Madejová, J., 2006. Acid activation of clay minerals. In: Bergaya, F., 1081
Theng, B. K. G., and Lagaly, G., Eds., Handbook of clay science. Elsevier. 1082
Lantenois, S., Champallier, R., Beny, J. M., and Muller, F., 2008. Hydrothermal synthesis and 1083
characterization of dioctahedral smectites: A montmorillonites series. Applied Clay 1084
Science 38, 165-178. 1085
45
Le Deit, L., Flahaut, J., Quantin, C., Hauber, E., Mege, D., Bourgeois, O., Gurgurewicz, J., 1086
Masse, M., and Jaumann, R., 2012. Extensive surface pedogenic alteration of the 1087
Martian Noachian crust suggested by plateau phyllosilicates around Valles Marineris. 1088
Journal of Geophysical Research-Planets 117. 1089
Loizeau, D., Mangold, N., Poulet, F., Bibring, J. P., Gendrin, A., Ansan, V., Gomez, C., 1090
Gondet, B., Langevin, Y., Masson, P., and Neukum, G., 2007. Phyllosilicates in the 1091
Mawrth Vallis region of Mars. Journal of Geophysical Research-Planets 112. 1092
Loizeau, D., Mangold, N., Poulet, F., Ansan, V., Hauber, E., Bibring, J. P., Gondet, B., 1093
Langevin, Y., Masson, P., and Neukum, G., 2010. Stratigraphy in the Mawrth Vallis 1094
region through OMEGA, HRSC color imagery and DTM. Icarus 205, 396-418. 1095
Longhi, J., Knittle, E., Holloway, J. R., and Wänke, H., 1992. The bulk composition, 1096
mineralogy and internal structure of Mars. In: Kieffer, H. H., Jakosky, B. M., Snyder, 1097
C. W., and Matthews, M. S., Eds., Mars. The University of Arizona Press, Tucson. 1098
Mangold, N., Poulet, F., Mustard, J. F., Bibring, J. P., Gondet, B., Langevin, Y., Ansan, V., 1099
Masson, P., Fassett, C., Head, J. W., Hoffmann, H., and Neukum, G., 2007. 1100
Mineralogy of the Nili Fossae region with OMEGA/Mars Express data: 2. Aqueous 1101
alteration of the crust. Journal of Geophysical Research-Planets 112. 1102
McLennan, S. M., 2003. Sedimentary silica on Mars. Geology 31, 315-318. 1103
McLennan, S. M., Anderson, R. B., Bell III, J. F., Bridges, J. C., Calef III, F., Campbell, J. L., 1104
Clark, B. C., Clegg, S., Conrad, P., Cousin, A., Des Marais, D. J., Dromart, G., Dyar, 1105
M. D., Edgar, L. A., Ehlmann, B. L., Fabre, C., Forni, O., Gasnault, O., Gellert, R., 1106
Gordon, S., Grant, J. A., Grotzinger, J. P., Gupta, S., Herkenhoff, K. E., Hurowitz, J. 1107
A., King, P. L., Le Mouélic, S., Leshin, L. A., Léveillé, R., Lewis, K. W., Mangold, 1108
N., Maurice, S., Ming, D. W., Morris, R. V., Nachon, M., Newsom, H. E., Ollila, A. 1109
M., Perrett, G. M., Rice, M. S., Schmidt, M. E., Schwenzer, S. P., Stack, K., Stolper, 1110
46
E. M., Sumner, D. Y., Treiman, A. H., VanBommel, S., Vaniman, D. T., Vasavada, 1111
A., Wiens, R. C., Yingst, R. A., and the MSL Science Team, 2014. Elemental 1112
geochemistry of sedimentary rocks at Yellowknife Bay, Gale crater, Mars. Science 1113
343, doi: 10.1126/science.1244734. 1114
Mitov, I., Paneva, D., and Kunev, B., 2002. Comparative study of the thermal decomposition 1115
of iron oxyhydroxides. Thermochimica Acta 386, 179-188. 1116
Murakami, T., Ito, J. I., Utsunomiya, S., Kasama, T., Kozai, N., and Ohnuki, T., 2004. Anoxic 1117
dissolution processes of biotite: implications for Fe behavior during Archean 1118
weathering. Earth and Planetary Science Letters 224, 117-129. 1119
Murchie, S. L., Mustard, J. F., Ehlmann, B. L., Milliken, R. E., Bishop, J. L., McKeown, N. 1120
K., Noe Dobrea, E. Z., Seelos, F. P., Buczkowski, D. L., Wiseman, S. M., Arvidson, 1121
R. E., Wray, J. J., Swayze, G., Clark, R. N., Des Marais, D. J., McEwen, A. S., and 1122
Bibring, J.-P., 2009. A synthesis of Martian aqueous mineralogy after one Mars year 1123
of observations from the Mars Reconnaissance Orbiter. Journal of Geophysical 1124
Research 114, E00D06. 1125
Mustard, J. F., Poulet, F., Gendrin, A., Bibring, J. P., Langevin, Y., Gondet, B., Mangold, N., 1126
Bellucci, G., and Altieri, F., 2005. Olivine and pyroxene diversity in the crust of Mars. 1127
Science 307, 1594-1597. 1128
Neubeck, A., Duc, N. T., Bastviken, D., Crill, P., and Holm, N. G., 2011. Formation of H(2) 1129
and CH(4) by weathering of olivine at temperatures between 30 and 70°C. 1130
Geochemical Transactions 12. 1131
Noe Dobrea, E. Z., Bishop, J. L., McKeown, N. K., Fu, R., Rossi, C. M., Michalski, J. R., 1132
Heinlein, C., Hanus, V., Poulet, F., Mustard, R. J. F., Murchie, S., McEwen, A. S., 1133
Swayze, G., Bibring, J. P., Malaret, E., and Hash, C., 2010. Mineralogy and 1134
stratigraphy of phyllosilicate-bearing and dark mantling units in the greater Mawrth 1135
47
Vallis/west Arabia Terra area: Constraints on geological origin. Journal of 1136
Geophysical Research-Planets 115. 1137
Ody, A., Poulet, F., Bibring, J. P., Loizeau, D., Carter, J., Gondet, B., and Langevin, Y., 2013. 1138
Global investigation of olivine on Mars: Insights into crust and mantle compositions. 1139
Journal of Geophysical Research: Planets 118, 234-262. 1140
Olsen, A. A. and Rimstidt, J. D., 2007. Using a mineral lifetime diagram to evaluate the 1141
persistence of olivine on Mars. American Mineralogist 92, 598-602. 1142
Owen, T., 1982. The composition of the Martian atmosphere. Advances in Space Research 2, 1143
75-80. 1144
Paris, M., Fritsch, E., and Reyes, B. O. A., 2007. H-1, Si-29 and Al-27 NMR study of the 1145
destabilization process of a paracrystalline opal from Mexico. Journal of Non-1146
Crystalline Solids 353, 1650-1656. 1147
Pokrovsky, O. S. and Schott, J., 2000. Kinetics and mechanism of forsterite dissolution at 25 1148
degrees C and pH from 1 to 12. Geochimica Et Cosmochimica Acta 64, 3313-3325. 1149
Poulet, F., Bibring, J. P., Mustard, J. F., Gendrin, A., Mangold, N., Langevin, Y., Arvidson, 1150
R. E., Gondet, B., and Gomez, C., 2005. Phyllosilicates on Mars and implications for 1151
early martian climate. Nature 438, 623-627. 1152
Poulet, F., Gomez, C., Bibring, J. P., Langevin, Y., Gondet, B., Pinet, P., Belluci, G., and 1153
Mustard, J., 2007. Martian surface mineralogy from Observatoire pour la Minéralogie, 1154
l'Eau, les Glaces et l'Activité on board the Mars Express spacecraft (OMEGA/MEx): 1155
Global mineral maps. Journal of Geophysical Research-Planets 112. 1156
Prasad, S. V. S. and Sitakara Rao, V., 1984. Thermal transformation of iron (III) oxide 1157
hydrate gel. Journal of Materials Science 19, 3266-3270. 1158
48
Rondeau, B., Fritsch, E., Guiraud, M., and Renac, C., 2004. Opals from Slovakia 1159
("Hungarian" opals): a re-assessment of the conditions of formation. European 1160
Journal of Mineralogy 16, 789-799. 1161
Santiago Buey, C., Barrios, M. S., Romero, E. G., and Montoya, M. D., 2000. Mg-rich 1162
smectite "precursor" phase in the Tagus Basin, Spain. Clays and Clay Minerals 48, 1163
366-373. 1164
Schröder, C., Klingelhöfer, G., and Tremel, M., 2004. Weathering of Fe-bearing minerals 1165
under Martian conditions, investigated by Mossbauer spectroscopy. Planetary and 1166
Space Science 52, 997-1010. 1167
Smith, K. L., Milnes, A. R., and Eggleton, R. A., 1987. Weathering of basalt: formation of 1168
iddingsite. Clays and Clay Minerals 35, 418-428. 1169
Tosca, N. J., Milliken, R. E., and Michel, F. M., 2008. Smectite formation on early Mars: 1170
experimental constraints. Martian phyllosilicates: recorders of aqueous processes, 1171
abstract #7030. 1172
Vaniman, D. T., Bish, D. L., Ming, D. W., Bristow, T. F., Morris, R. V., Blake, D. F., 1173
Chipera, S. J., Morrison, S. M., Treiman, A. H., Rampe, E. B., Rice, M., Achilles, C. 1174
N., Grotzinger, J., McLennan, S. M., Williams, J., Bell III, J., Newsom, H., Downs, R. 1175
T., Maurice, S., Sarrazin, P., Yen, A. S., Morookian, J. M., Farmer, J. D., Stack, K., 1176
Milliken, R. E., Ehlmann, B., Sumner, D. Y., Berger, G., Crisp, J. A., Hurowitz, J. A., 1177
Anderson, R., DesMarais, D., Stolper, E. M., Edgett, K. S., Gupta, S., Spanovich, N. 1178
and the MSL Science Team, 2014. Mineralogy of a mudstone at Yellowknife Bay, 1179
Gale crater, Mars. Science 343, doi: 10.1126/science.1243480. 1180
Velbel, M. A., Long, D. T., and Gooding, J. L., 1991. Terrestrial weathering of Antarctic 1181
stone meteorites: Formation of Mg-carbonates on ordinary chondrites. Geochimica et 1182
Cosmochimica Acta 55, 67-76. 1183
49
Wilson, M. J., 2004. Weathering of the primary rock-forming minerals: processes, products 1184
and rates. Clay Minerals 39, 233-266. 1185
Wogelius, R. A. and Walther, J. V., 1992. Olivine dissolution kinetics at near-surface 1186
conditions. Chemical Geology 97, 101-112. 1187
Wordsworth, R., Forget, F., Millour, E., Head, J. W., Madeleine, J. B., and Charnay, B., 2012. 1188
Global modelling of the early martian climate under a denser CO2 atmosphere: Water 1189
cycle and ice evolution. Icarus 222, 1-19. 1190
Wray, J. J., Murchie, S. L., Squyres, S. W., Seelos, F. P., and Tornabene, L. L., 2009. Diverse 1191
aqueous environments on ancient Mars revealed in the southern highlands. Geology 1192
37, 1043-1046. 1193
Yen, A. S., Kim, S. S., Hecht, M. H., Frant, M. S., and Murray, B., 2000. Evidence that the 1194
reactivity of the martian soil is due to superoxide ions. Science 289, 1909-1912. 1195
Zent, A. P., 1998. On the thickness of the oxidized layer of the Martian regolith. Journal of 1196
Geophysical Research-Planets 103, 31491-31498. 1197
1198
50
FIGURES AND TABLES 1199
1200
Fig. 1. NIR spectrum of the initial olivine used in the experiment. Left panel: general aspect 1201
of the spectrum, exhibiting a typical olivine signature with a broad Fe2+-related absorption 1202
band at 1.04 µm. Left panel, inset: small apatite crystals (arrows) observed by SEM on the 1203
surface of some olivine grains. Middle panel: detail of the spectrum of the initial olivine, 1204
revealing a very shallow band at 2.31 µm (top), compared to a reference spectrum of apatite 1205
(bottom; sample LAAP03 of CRISM spectral library). Note the different reflectance scales. 1206
Right panel: another detail of the spectrum, revealing a small absorption band at 4.02 µm 1207
(top), which may also be due to apatite despite the slight shift in comparison with the 1208
reference spectrum (bottom). 1209
1210
Fig. 2. Schematic diagram and photograph of the experimental device, and summary of the 1211
weathering conditions tested. 1212
1213
Fig. 3. Chemical evolution with time of the experimental solutions. Left panel: evolution of 1214
pH (diamonds). Theoretical starting values (triangles) were calculated using the JCHESS 1215
geochemical software for pure water under CO2 partial pressures of 1 (CO2 atmosphere) and 1216
4.10-4 (terrestrial atmosphere). Error bars are added for unstable measures. Right panel: 1217
evolution of Si, Mg and Fe mass concentrations. Error bars include analytical uncertainties 1218
(provided by the SARM laboratory) and further uncertainties related to contaminations by the 1219
experimental device itself, which were determined thanks to a blank experiment (without 1220
solid phase). See section 2.4.4 for details about the protocols. 1221
1222
51
Fig. 4. NIR spectra of the final (weathered) samples of the experiment, compared to the 1223
spectrum of the initial olivine and to some library spectra. A: general aspect of the spectra 1224
between 1.0 and 2.6 µm. B: detailed view between 1.8 and 2.5 µm. C: detailed view between 1225
3.8 and 4.2 µm. Arrows in B and C indicate newly-appeared or deepened absorption bands. 1226
D: reference spectra of clay minerals and opal from the USGS library (Clark et al., 2007) 1227
(nontronite NG-1a, saponite SapCa-1, montmorillonite CM20; opal TM8896-hyalite). E: 1228
spectrum of a natural sample of nesquehonite (mixed with other Mg-carbonates) (Dehouck et 1229
al., 2012). F: reference spectra of carbonate minerals from the USGS library (hydromagnesite 1230
LACB28A, magnesite LACB03A, siderite LACB08A). 1231
1232
Fig. 5. Band depths calculations for the initial and weathered olivine powders. Error bars 1233
correspond to standard deviation around the average (diamonds) for six different 1234
measurements. For the 1.91-µm band, the continuum anchors are taken at 1.80 and 2.10 µm. 1235
For the 2.31-µm band, the continuum anchors are taken at 2.25 and 2.36 µm. In both cases, 1236
every values used in the calculation correspond to the median of three spectral channels to 1237
minimize the effect of instrumental noise. 1238
1239
Fig. 6. NIR spectra of the <2-µm fraction of the weathered olivine powders. The slight shift 1240
of the Fe2+-related band at 1.07 µm compared to Fig. 4 is an effect of the lower grain size. The 1241
detailed, continuum-corrected view on the right highlights the 2.31-µm absorption band, 1242
whose depth varies between samples. 1243
1244
Fig. 7. X-ray diffraction patterns (Cu Kα radiation) obtained from the oriented <2-µm fraction 1245
of the initial and weathered olivine powders. No diffraction peaks corresponding to clay 1246
minerals are observed in any sample. 1247
52
Fig. 8. TEM micrographs showing the different textures observed at the surface of the olivine 1248
grains (in dark) before and after the experiment. A: examples of clean surfaces resulting from 1249
grinding and typical of the initial olivine. B: examples of “cotton-like” textures, observed 1250
only in the “CO2” and “CO2-H2O2” samples. Note a nearby grain with clean surfaces in the 1251
top image. C: examples of filamentous textures, observed in the “Air”, “Air-H2O2” and 1252
“CO2” samples. D: comparison of the filamentous texture found in this study (left) with 1253
phyllosilicates (serpentine and saponite) produced by experimental alteration of olivine at 1254
200°C (Jones and Brearley, 2006) (right). The two micrographs are presented at the same 1255
scale. 1256
1257
Fig. 9. Chemical compositions measured by TEM-EDX in the initial and weathered olivine 1258
samples. A: ternary diagram Si-Mg-Fe showing individual compositions measured on initial 1259
grains and newly-formed phases (see caption). B to E: same data as in A, redrawn in the form 1260
of separate color ranges for clarity. 1261
1262
Fig. 10. Differential thermogravimetric (DTG) curves of the initial and weathered olivine 1263
powders. (“Air-H2O2” was not analyzed, having similar properties than “Air” in other 1264
datasets.) Samples were heated at 10 °C/min in dinitrogen. Mass losses are indicated by 1265
arrows and interpreted in terms of volatile releases (ads.: adsorbed). The mass loss at 760 °C 1266
in the curve of the “CO2” sample is probably a noise artifact. 1267
1268
Table 1. Chemical composition of the initial olivine used in the experiment, as determined by 1269
ICP-OES. LOD: limit of detection. LOI: loss on ignition (measured by thermogravimetry and 1270
heating at 1000 °C). 1271
1272
53
Table 2. Summary of analytical results and interpretations in terms of secondary phases. For 1273
NIR spectroscopy and TEM observations, the number of “+” symbols indicates the relative 1274
intensity of the absorption bands and the relative abundances of the newly-formed phases as 1275
estimated by eye, respectively. aEstimation based on the amounts of H2O and OH released 1276
from the corresponding samples, assuming a water content of 6 to 13 wt% for the Si-rich 1277
phase (Knauth and Epstein, 1982; Paris et al., 2007). The value may be slightly overestimated 1278
for the “CO2” sample, for which some H2O is borne by the smectite phase. bUpper limit of ~1 1279
wt% deduced from the lower abundance and thickness of the smectite/filamentous phase 1280
compared to the Si-rich/cotton-like phase, as observed by TEM. In addition, NIR and TEM 1281
indicate that the smectite phase is less abundant in the “CO2” sample than in the two “Earth 1282
samples”. cEstimation based on the amount of CO2 released in thermogravimetric analyses 1283
(attributed to magnesite and siderite, see text). dUpper limit estimated from the amount of 1284
dissolved Fe measured in the “CO2” reactor (24 mg/L at 14 days; Table S1), assuming that all 1285
of this Fe is precipitated as (oxy)hydroxide. ads.: adsorbed. n. a.: not analyzed. LOD: limit of 1286
detection.1287
54
1288
Fig. 1. NIR spectrum of the initial olivine used in the experiment. Left panel: general aspect 1289
of the spectrum, exhibiting a typical olivine signature with a broad Fe2+-related absorption 1290
band at 1.04 µm. Left panel, inset: small apatite crystals (arrows) observed by SEM on the 1291
surface of some olivine grains. Middle panel: detail of the spectrum of the initial olivine, 1292
revealing a very shallow band at 2.31 µm (top), compared to a reference spectrum of apatite 1293
(bottom; sample LAAP03 of CRISM spectral library). Note the different reflectance scales. 1294
Right panel: another detail of the spectrum, revealing a small absorption band at 4.02 µm 1295
(top), which may also be due to apatite despite the slight shift in comparison with the 1296
reference spectrum (bottom). 1297
1298
55
1299
Fig. 2. Schematic diagram and photograph of the experimental device, and summary of the 1300
weathering conditions tested. 1301
1302
56
1303
Fig. 3. Chemical evolution with time of the experimental solutions. Left panel: evolution of 1304
pH (diamonds). Theoretical starting values (triangles) were calculated using the JCHESS 1305
geochemical software for pure water under CO2 partial pressures of 1 (CO2 atmosphere) and 1306
4.10-4 (terrestrial atmosphere). Error bars are added for unstable measures. Right panel: 1307
evolution of Si, Mg and Fe mass concentrations. Error bars include analytical uncertainties 1308
(provided by the SARM laboratory) and further uncertainties related to contaminations by the 1309
experimental device itself, which were determined thanks to a blank experiment (without 1310
solid phase). See section 2.4.4 for details about the protocols. 1311
1312
57
1313
Fig. 4. NIR spectra of the final (weathered) samples of the experiment, compared to the 1314
spectrum of the initial olivine and to some library spectra. A: general aspect of the spectra 1315
between 1.0 and 2.6 µm. B: detailed view between 1.8 and 2.5 µm. C: detailed view between 1316
3.8 and 4.2 µm. Arrows in B and C indicate newly-appeared or deepened absorption bands. 1317
58
D: reference spectra of clay minerals and opal from the USGS library (Clark et al., 2007) 1318
(nontronite NG-1a, saponite SapCa-1, montmorillonite CM20; opal TM8896-hyalite). E: 1319
spectrum of a natural sample of nesquehonite (mixed with other Mg-carbonates) (Dehouck et 1320
al., 2012). F: reference spectra of carbonate minerals from the USGS library (hydromagnesite 1321
LACB28A, magnesite LACB03A, siderite LACB08A). 1322
1323
59
1324
Fig. 5. Band depths calculations for the initial and weathered olivine powders. Error bars 1325
correspond to standard deviation around the average (diamonds) for six different 1326
measurements. For the 1.91-µm band, the continuum anchors are taken at 1.80 and 2.10 µm. 1327
For the 2.31-µm band, the continuum anchors are taken at 2.25 and 2.36 µm. In both cases, 1328
every values used in the calculation correspond to the median of three spectral channels to 1329
minimize the effect of instrumental noise. 1330
1331
60
1332
Fig. 6. NIR spectra of the <2-µm fraction of the weathered olivine powders. The slight shift 1333
of the Fe2+-related band at 1.07 µm compared to Fig. 4 is an effect of the lower grain size. The 1334
detailed, continuum-corrected view on the right highlights the 2.31-µm absorption band, 1335
whose depth varies between samples. 1336
1337
61
1338
Fig. 7. X-ray diffraction patterns (Cu Kα radiation) obtained from the oriented <2-µm fraction 1339
of the initial and weathered olivine powders. No diffraction peaks corresponding to clay 1340
minerals are observed in any sample. 1341
1342
62
1343
Fig. 8. TEM micrographs showing the different textures observed at the surface of the olivine 1344
grains (in dark) before and after the experiment. A: examples of clean surfaces resulting from 1345
grinding and typical of the initial olivine. B: examples of “cotton-like” textures, observed 1346
only in the “CO2” and “CO2-H2O2” samples. Note a nearby grain with clean surfaces in the 1347
top image. C: examples of filamentous textures, observed in the “Air”, “Air-H2O2” and 1348
“CO2” samples. D: comparison of the filamentous texture found in this study (left) with 1349
phyllosilicates (serpentine and saponite) produced by experimental alteration of olivine at 1350
200°C (Jones and Brearley, 2006) (right). The two micrographs are presented at the same 1351
scale.1352
63
1353
Fig. 9. Chemical compositions measured by TEM-EDX in the initial and weathered olivine 1354
samples. A: ternary diagram Si-Mg-Fe showing individual compositions measured on initial 1355
grains and newly-formed phases (see caption). B to E: same data as in A, redrawn in the form 1356
of separate color ranges for clarity. 1357
1358
64
1359
Fig. 10. Differential thermogravimetric (DTG) curves of the initial and weathered olivine 1360
powders. (“Air-H2O2” was not analyzed, having similar properties than “Air” in other 1361
datasets.) Samples were heated at 10 °C/min in dinitrogen. Mass losses are indicated by 1362
arrows and interpreted in terms of volatile releases (ads.: adsorbed). The mass loss at 760 °C 1363
in the curve of the “CO2” sample is probably a noise artifact. 1364
1365
1366
65
1367
wt.%
SiO2 40.50
Al2O3 0.08
Fe2O3 10.44
MnO 0.13
MgO 50.09
CaO 0.10
Na2O 0.02
K2O <LOD
TiO2 <LOD
P2O5 <LOD
LOI -0.74
Total 100.63
1368
Table 1. Chemical composition of the initial olivine used in the experiment, as determined by 1369
ICP-OES. LOD: limit of detection. LOI: loss on ignition (measured by thermogravimetry and 1370
heating at 1000 °C). 1371
1372
66
1373
CO2-H2O2 CO2 Air-H2O2 Air
Analytical results
NIR bands 1.91 µm ++ ++ ++ ++
2.31 µm + ++ ++
TEM Cotton-like + ++
Filamentous + ++ ++
DTG mass
losses (%)
ads. H2O ~0.09 ~0.10 n. a. ~0.07
OH ~0.06 ~0.15 n. a.
CO2 ~0.02 n. a.
TIC (mg/g) <LOD 0.09 n. a. <LOD
Secondary phases (wt%)
Si-rich phasea ~1 – 3 ~2 – 4
Smectiteb <<1 <1 <1
Carbonatesc ~0.05
Fe-(oxy)hydroxidesd <0.1 <0.1
1374
Table 2. Summary of analytical results and interpretations in terms of secondary phases. For 1375
NIR spectroscopy and TEM observations, the number of “+” symbols indicates the relative 1376
intensity of the absorption bands and the relative abundances of the newly-formed phases as 1377
estimated by eye, respectively. aEstimation based on the amounts of H2O and OH released 1378
from the corresponding samples, assuming a water content of 6 to 13 wt% for the Si-rich 1379
phase (Knauth and Epstein, 1982; Paris et al., 2007). The value may be slightly overestimated 1380
for the “CO2” sample, for which some H2O is borne by the smectite phase. bUpper limit of ~1 1381
wt% deduced from the lower abundance and thickness of the smectite/filamentous phase 1382
compared to the Si-rich/cotton-like phase, as observed by TEM. In addition, NIR and TEM 1383
indicate that the smectite phase is less abundant in the “CO2” sample than in the two “Earth 1384
samples”. cEstimation based on the amount of CO2 released in thermogravimetric analyses 1385
(attributed to magnesite and siderite, see text). dUpper limit estimated from the amount of 1386
dissolved Fe measured in the “CO2” reactor (24 mg/L at 14 days; Table S1), assuming that all 1387
of this Fe is precipitated as (oxy)hydroxide. ads.: adsorbed. n. a.: not analyzed. LOD: limit of 1388
detection. 1389
