Evaluation of carbonate abundance in putative martian paleolake basins

Icarus 200 (2009) 426–435

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Evaluation of carbonate abundance in putative martian paleolake basins V. Orofino a,∗ , J. Goldspiel b , I. Carofalo a , A. Blanco a , S. Fonti a , G.A. Marzo a a b

Dipartimento di Fisica, Università del Salento, Via Arnesano, C.P. 193, 73100 Lecce, Italy US Naval Research Laboratory, Space Science Division, 4555 Overlook Avenue SW, Washington, DC 20375, USA

a r t i c l e

i n f o

a b s t r a c t

Article history: Received 12 November 2007 Revised 11 November 2008 Accepted 20 November 2008 Available online 11 December 2008 Keywords: Mars, surface Geological processes

Carbonate deposits have not been found so far on Mars, although there appears to have been sufficient water to have supported their formation. Many hypotheses have been proposed in order to explain this. In the present work we explore the possibility that the missed detection of carbonate deposits on the martian surface could be simply due to the fact that the concentration of carbonates, when mixed with other materials present in the sedimentary deposits, may be below the detection limit of the various instruments used so far in this search. In the present study we consider 21 putative paleolacustrine basins and use a sediment transport model to estimate the abundance of carbonates which could be present in the sediments deposited on the basin floor. In this way we find that for all the selected basins the estimated carbonate abundances are in general less than a few percent, and such values are below (or at best comparable to) the detection limits of the spectrometers flown around Mars during the recent space missions. Furthermore, applying the sediment transport model to the well studied Eberswalde crater, we conclude that the fluvio-lacustrine activity in this basin should have lasted for a period on the order of 103 –104 years, in good agreement with previous work. Our results suggest that a hydrological cycle, able to move large volumes of water and to create relatively stable lakes, could have been active intermittently on Mars in the past, producing carbonate deposits that could escape detection by the instruments that have flown to date. © 2008 Elsevier Inc. All rights reserved.

1. Introduction Today on the surface of Mars the average temperature, close to 210 K, and the atmospheric pressure (about 6 mbar) are low enough to prevent the stable presence of liquid water, which usually freezes solid and then quite rapidly sublimates, except in the polar regions where water ice is relatively stable. Furthermore, due the very low water content in the atmosphere, any form of precipitation, whether rain or snow, is practically impossible. However, it is likely that in the past the planet may have experienced episodes of warmer and wetter climate with a perhaps thicker atmosphere able to stabilize liquid water on the martian surface (Pollack et al., 1987), even if it is not understood how a warm and wet climate could have been sustained on early Mars (Kasting, 1991; Haberle, 1998). This scenario is suggested by the presence of a large number of interesting geologic features on the martian surface. For example dendritic valley networks, apparently carved by runoff (that is, through slow erosion by water running across the surface), are common in the southern highlands of the planet (Masursky

*

Corresponding author. E-mail address: vincenzo.orofi[email protected] (V. Orofino).

0019-1035/$ – see front matter doi:10.1016/j.icarus.2008.11.020

©

2008 Elsevier Inc. All rights reserved.

et al., 1977; Malin and Carr, 1999), and some of them have attributes suggestive of an origin due to precipitation (Craddock and Howard, 2002; Mangold et al., 2004; Ansan and Mangold, 2006). In addition, many valleys often debouch into basins, such as depressions or impact craters, which possibly hosted lakes in the past (Goldspiel and Squyres, 1991; Forsythe and Blackwelder, 1998; Cabrol and Grin, 1999). This is also suggested by the presence of massive layered outcrops in the interiors of craters and chasmas, first recognized in Valles Marineris and interpreted as sedimentary deposits (McCauley, 1978; Malin and Edgett, 2000). Interestingly, spectroscopic observations of these layered outcrops, made by the spectrometer OMEGA (Observatoire pour la Minéralogie, l’Eau, la Glace et l’Activité) onboard the Mars Express spacecraft, have shown the presence of magnesium and calcium sulfates (Bibring et al., 2005; Gendrin et al., 2005), which, by analogy with terrestrial geology, could originate from evaporation of standing bodies of water such as shallow lakes or lagoons (Warren, 1999), even if alternative hypotheses have been proposed (Knauth et al., 2005). In Terra Meridiani, light-toned sulfate-rich layers coexist with coarse-grained crystalline hematite deposits, discovered by the Thermal Emission Spectrometer (TES) onboard the Mars Global Surveyor spacecraft (Christensen et al., 2000). This coincidence, also detected in situ by the Mars Exploration Rover Opportunity (Squyres et al., 2004a; Glotch et al.,

Carbonate abundance in putative martian paleolakes

2006), is intriguing since the formation of coarse-grained crystalline hematite (Fe2 O3 ) could be due to the presence of an ironrich aqueous medium (Christensen et al., 2000). Furthermore, not far from the Opportunity landing site, as well as in other areas near the boundary of the hemispheric dichotomy, other indicators, such as paleoshoreline morphologies, would suggest the past existence of an ocean (Parker et al., 1989, 1993; Baker et al., 1991; Webb, 2004), although this interpretation remains quite controversial (Malin and Edgett, 1999; Ghatan et al., 2006). In any case, by analogy with terrestrial examples, these geologic features suggest the presence of liquid water on the surface for a significant period of time. In this context, the detection on Mars of evaporitic minerals, such as carbonates, sulfates and other salts, could play a crucial role in understanding the climatic history of the planet. Actually, as reported above, magnesium and calcium sulfates have been identified by Mars Express and Opportunity in various small areas of the planet (Bibring et al., 2005). In spite of the limited extent of these deposits, this discovery is important since the martian sulfates could in principle be present within a much more extended subsurface layer, appearing at the surface only in the few places detected by OMEGA. On the other hand, apart from the identification of small quantities of carbonates in the martian dust (Pollack et al., 1990; Lellouch et al., 2000; Bandfield et al., 2003), no carbonate deposits have been found so far on Mars (McKay and Nedell, 1988; Christensen et al., 2001; Bibring et al., 2005; Stockstill et al., 2005, 2007; Jouglet et al., 2007). In addition to the obvious assumption that massive carbonate deposits never formed on Mars (Bullock and Moore, 2007), many other hypotheses have been proposed in order to explain why remote sensing instruments have not identified such deposits. For a detailed discussion on the subject see Craddock and Howard (2002). Here we explore the possibility that the missed detection of carbonate deposits on the martian surface could be simply due to the fact that the concentration of carbonates, when mixed with the other sedimentary materials, may be lower than the detection limit of the various instruments used so far in this kind of search (Hartmann et al., 2001). For this study we consider a sample of putative paleolacustrine basins and first use a sediment transport model (Bagnold, 1966) to estimate the volume of water necessary to remove the sediments and carve the inflow valleys. Then, with reasonable assumptions on the content of calcium and magnesium ions in martian groundwaters, we evaluate the abundance of carbonates in the sediments present on the basin floor. In Section 2, we present a general description of the approach followed in this paper to obtain the expected carbonate abundances in the sample basins; the assumptions that are at the basis of our method are also discussed. In Section 3, we describe the criteria used to select the basins and the procedure to evaluate the sediment volumes. In Section 4, we discuss in detail the method to calculate the water volumes and the associated carbonate abundances for the chosen sites and we present our results, comparing them with previous similar evaluations. In Section 5, we report the results concerning the estimated carbonate abundances in the selected basins. These abundances are generally below (or at best comparable to) the detection limits of the four spectrometers flown on Mars Global Surveyor, Mars Express and Mars Reconnaissance Orbiter, if the carbonate precipitates are homogeneously mixed with the clastic sediments transported into the lake basins by the inflow streams. This may explain why none of these instruments have detected carbonate deposits on Mars. In Section 6, we apply Bagnold’s sediment transport model to the well studied Eberswalde crater, in order to evaluate the duration of the fluviolacustrine activity in this basin and compare our estimates for this

427

case with those obtained by other authors. Finally in Section 7, we discuss our results and report the conclusions of our work. 2. Methodological approach The most important part of our work is the estimate of the volumes of water necessary to carve the valleys that debouch into craters that acted in the past as lacustrine basins. This can be done, by means of a sediment transport model (Bagnold, 1966), starting from the previously determined sediment volumes removed and transported by ancient rivers once present on the floor of the inflow valleys. The Bagnold model provides sediment transport rate solutions for bedload and suspended load that, because of their basis in general physics principles, are readily adaptable to Mars applications (Komar, 1979; Goldspiel and Squyres, 1991). This model is intended for fully turbulent streams in which the thickness of the bedload zone is negligible in comparison to the water height. As we will discuss in Section 7, these conditions are met in the cases considered here. In calculating the total water volume in this manner, we assume that no work is done by the stream to break down the regolith into sediment grains, and so the water volume calculations are conservative estimates for the amount of water required to carve the valleys. Another implicit assumption of the model is that the studied craters acted in the past as accumulation basins for all the sediments transported by the flow, so that the whole mass of the materials eroded from the valleys is still present inside the basins. For this reason the model can be applied to so-called “closed” craters (Forsythe and Blackwelder, 1998; Cabrol and Grin, 1999), that is impact structures associated with one or more inflow valleys but for which no spillways are observed. In order to obtain the volume of water involved in the erosion of the inlet valleys, the sediment size distribution has to be known. In this respect, we follow two different approaches, assuming that: (a) the sediments we see today on the floor of the craters are the same as those deposited by the inlet channels during the past periods of fluvio-lacustrine activity; (b) the sediments of all the basins have the same size distribution as that determined for the sedimentary environment of Meridiani Planum by the Mars Exploration Rover Microscopic Imager at the landing site of Opportunity rover (Squyres et al., 2004b). In the former case, we adopt the optimistic hypothesis that the basins that are likely sites of aqueous sedimentation were not subsequently filled with lava or aeolian materials which would bury the potential lacustrine deposits and prevent their detection by remote sensing spectroscopic observations. Even if there are clear cases of volcanic resurfacing within craters (and Gusev crater is one of the most important examples; Martínez-Alonso et al., 2005), our hypothesis is not unreasonable. For all the chosen sites in this study, there is no convincing evidence for volcanic resurfacing (Cabrol and Grin, 1999). Once the water volumes that may have flowed into the basins are known, for the chosen sites the carbonate content can be obtained, following the same approach adopted by Goldspiel and Squyres (1991), by means of simple stoichiometric calculations that use as input parameters the assumed concentrations of Mg2+ and Ca2+ ions in the martian waters. In conclusion, our approach can be divided in four steps: 1) 2) 3) 4)

selection of the sites; evaluation of the sediment volumes; estimation of the water volumes; calculation of the carbonate abundances.

These steps will be described in detail in the next two sections.

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3. Site selection and sediment volume evaluation Traces of ancient lacustrine and fluvial activity are widespread on the martian surface, although they occur under different conditions and their interpretation is at places debatable. In particular, lakes in martian impact craters have been the subject of several local and regional studies (see Cabrol and Grin, 1999, and references therein). Martian paleolakes have been, in fact, suggested as landing sites for in situ and sample-return missions since they should provide information about the dynamics of the sedimentary processes and the climate under which they were formed. They are also interesting since they can be regarded as potential sites for the detection of biomarkers. In order to select the sites suitable for our calculations of carbonate abundances, we have looked at sites where geologic features (terraces, deposits, inlet and/or outlet channels and putative deltas) suggest the ancient presence of standing bodies of water. We started with a sample of 222 sites that Orofino et al. (2004) considered suitable for a spectroscopic search for carbonates by means of the Planetary Fourier Spectrometer (PFS) instrument onboard the Mars Express spacecraft. The sites were selected by merging three previous catalogues of proposed paleolacustrine basins published by different authors, namely: 144 closed crater basins analyzed by Forsythe and Blackwelder (1998), 179 different impact crater lakes (closed, open and lake-chain systems) studied by Cabrol and Grin (1999), and 45 basins (both craters and depressions) extracted from several monographic sources by Orofino et al. (2000). In our sample we have not included the large basins (like Elysium, Amazonis and others) that, according to various authors (see Scott et al., 1995, and references therein), have harbored wide lakes or seas. In fact, according to Craddock and Maxwell (1993), the formation of carbonate deposits is more likely in relatively small pools of water, due to the high concentration of the eroded materials which raises the local pH of the water and favors the deposition of carbonates. For the reasons discussed in the previous section, we initially extracted 167 sites, out of the starting 222, that are listed as “closed basins” in the above cited catalogues. After this first selection, using an interactive Viking data map, we performed a thorough analysis of the morphological properties of the chosen sites, in order to select only the systems that appear well-developed with regular and well defined edges. At this stage, in fact, we excluded all craters and channels characterized by uneven ground and jagged boundaries, as well as those that appear to have undergone subsequent impacts. In addition, we considered only craters that are at least 20 km across and have one or more inlet valleys longer than the crater radius. This last selection method was adopted in order to exclude all the basins with very short inlet channels, which could mean that in the past only small amounts of water and sediment flowed into the basins and that the basin surface was only intermittently covered by a thin layer of water. Alternatively, the influx of water necessary to feed this kind of lake could have derived from direct precipitation (Parker, 1997) or, more likely, through groundwater leakage into the basin (Carr, 1996). In both cases our approach is less useful in evaluating the water volumes involved, and the corresponding carbonate abundances, for basins with short valleys. For this reason, these small systems were excluded from our sample list. With these very restrictive selection criteria, we extracted 28 useful sites to be further analyzed using Mars Orbiter Laser Altimeter (MOLA) topographic data. The MOLA data were used to evaluate the longitudinal profile of the ground near each inflow valley and to create the topographic cross-section profiles at different locations along the valleys; this method allowed us to assess the valley slopes, which are in the range 0.006–0.05 (see Table 1),

Table 1 Volume of sediments eroded from all the inflow valleys of the sample basins and slope of the valleys. For basins with more than one main inlet valley, the range of valley slopes is given. The crater No. 7 (the only one to have an official IAU name) is called Tuscaloosa crater. Site No.

Crater coordinates

Eroded sediment volume (km3 )

Slope

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20

35.0◦ N; 8.0◦ W 22.0◦ N; 306.0◦ W 10.3◦ N; 16.7◦ W 8.5◦ N; 56.5◦ W 6.5◦ N; 321.2◦ W 5.0◦ N; 58.6◦ W 0.0; 331.0◦ W 5.7◦ S; 331.2◦ W 6.0◦ S; 307.2◦ W 9.5◦ S; 261.3◦ W 11.8◦ S; 347.5◦ W 12.5◦ S; 299.5◦ W 15.5◦ S; 298.5◦ W 16.9◦ S; 295.0◦ W 18.2◦ S; 228.4◦ W 18.8◦ S; 185.2◦ W 19.2◦ S; 300.8◦ W 19.6◦ S; 307.9◦ W 30.2◦ S; 188.5◦ W 36.6◦ S; 72.7◦ W

2.0 46.0 23.0 1.0 23.0 18.0 25.0 13.5 20.4 5.0 9.4 96.0 3.4 34.0 1.0 45.0 46.5 5.8 2.0 4.0

0.01 0.01 0.01 0.01 0.02 0.01 0.003 0.01 0.006–0.01 0.007 0.02 0.01–0.03 0.01–0.02 0.03–0.04 0.01 0.008–0.010 0.02–0.05 0.01 0.006 0.03

and also to calculate the inlet valley volumes that approximate the sediment volumes eroded and transported by water into each crater (see below). It is important to note that in some sites the slope of the ground around the basins appears irregular and uncertain, so it is difficult to assess if the associated channels actually were lake outlet or inlet channels that subsequently experienced some process that modified their morphological characteristics. Indeed the analysis of MOLA data showed that 10 basins among the selected sites listed in the literature as “closed” are in fact “open.” In other words, some of the channels associated with these craters, considered on the basis of Viking data as tributaries, should actually have played the role of outlets. On the other hand, the two craters at 12.5◦ S/299.5◦ W (Site No. 12 in Table 1) and at 18.8◦ S/185.2◦ W (Site No. 16), reported in the Cabrol and Grin (1999) catalogue as “open systems,” should actually be considered closed craters because they exhibit sharp inlet valley systems, while the structures which could be spillways are not so sharp and well developed. In fact they are present in the sample of closed craters analyzed by Forsythe and Blackwelder (1998). The total number of sites suitable for our analysis, then, is 20. Two of these sites are shown in Figs. 1 and 2. For each site, the total volume of the eroded material has been estimated from all of the inflow valleys, assuming that all the material was deposited in the crater as sediment. Using the method of Goldspiel and Squyres (1991), each valley was partitioned into a series of contiguous segments; the segments were treated as trapezoidal prisms, representing on average the geometric shape of the analyzed channels, so that each prism shared its two bases with the adjoining sections. For each valley segment, the volume was calculated by multiplying the average cross-section area included between the floor and the top of valley walls by the length of the valley segment. The cross-section areas were estimated from MOLA data profiles (available at the web site http://marsoweb.nas.nasa.gov), while the lengths of the prisms were evaluated using the interactive Mars maps available at the USGS web site (http://webgis.wr.usgs.gov/ website/mars_html/viewer.htm). The volumes of the individual valley segments were then summed to approximate the total sediment volume of the main trunk. An analogous approach was followed for the tributaries. In this way we evaluated, for each basin,

Carbonate abundance in putative martian paleolakes

429

Fig. 1. Study site No. 12. Example of a potential paleolacustrine basin, about 44 km across, with a relatively well-developed inflow valley system. The basin is located at 12.5◦ S, 299.5◦ W. The image is taken from THEMIS database by means of JMARS software (at http://jmars.asu.edu/).

Vw =

Fig. 2. Study site No. 20. Example of a potential paleolacustrine basin, 33 km across, with a basic inflow valley system. The basin is located at 36.6◦ S, 72.7◦ W. The image is taken from THEMIS database by means of JMARS software (at http://jmars. asu.edu/).

the total sediment volume eroded from all the inflow valleys (tributaries included), as reported in Table 1. 4. Water volumes and carbonate abundances With the Bagnold’s sediment transport rate, the total volume V w of water required to transport sediment of volume V s is given by (Goldspiel and Squyres, 1991):

u¯ H (ρs − ρ w ) g i

V s,

(1)

where i is the sediment transport rate in units of immersed weight per unit stream width per second, u¯ is the mean stream flow velocity calculated from the stream channel geometry and a dimensionless drag coefficient, H is the channel hydraulic radius (open channel stream depth), ρs and ρ w are the sediment and water densities, while g is the gravitational acceleration at the surface of Mars. In our case the above quantities refer to the interior channel once present on the floor of each valley that debouches in the basin under study. As far as the hydraulic radius H is concerned, it cannot be directly measured from the available images of our sample basins, since no interior channels are visible in the inflow valleys associated with the basins. Actually, interior channels are rare on Mars, since they have been observed only within an unnamed valley in Libya Montes (Jaumann et al., 2005) and in 21 other locations studied by Irwin et al. (2005). For this reason the values of H have to be assumed; in particular we take H = 1 m and H = 5 m, consistent with the range typically observed in small terrestrial rivers. These values are also consistent with the geometry of the interior channel located in Libya Montes, studied by Jaumann et al. (2005). For this channel, the authors measure a width w = 450 m and a depth D = 40 m, evaluating a hydraulic radius H = 3 m (for a 10% bankfull discharge of about 4800 m3 s−1 ). For comparison, Williams (1988) finds that, for terrestrial rivers, width and hydraulic radius of the streams are linked by the empirical law H = ξ w 0.69 , where ξ is a constant and H and w are in meters. Assuming as a rough approximation that the same law holds for martian streams, the interior channel in Libya Montes allows us to calibrate the relationship between width and hydraulic radius, producing a value of 0.044 (in MKS units) for the constant ξ . Applying this power law to the 21 interior channels studied by Irwin et al. (2005), for which 90 m  w  1000 m, we obtain 1.0 m  H  5.2 m, in excellent agreement with our choices. Following the approach by Goldspiel and Squyres (1991), it is possible to evaluate the other two quantities i and u¯ present in Eq. (1), if the mean grain size and the slope of the channel are

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known. As far as the quantity i is concerned, in Bagnold’s theory it is linked to an empirical parameter called “bedload efficiency factor,” eb . In his pioneering work Bagnold (1966) provides the values of this parameter for a set of only four values of mean grain diameter (0.03, 0.1, 0.3 and 1 mm). So we have to interpolate and partially extrapolate the values of eb in order to have data in the range of 0.01–3.0 mm. We adopt this range because it encompasses silt, sand and small pebbles common in terrestrial river sediments (e.g., Bagnold, 1966), and also because outside this range a significant extrapolation of Bagnold’s experimental data would be required, leading to highly uncertain results. As a first step, for each site, the mean grain diameter of the regolith is calculated starting from thermal inertia measurements. In fact, the thermal inertia I of the martian soil is linked the thermal conductivity k by the equation (Carr, 1981): I=



kρ c ,

(2)

where ρ is the bulk density and c is specific heat of the regolith. Once the thermal conductivity is known, the mean grain size d of the martian regolith can be derived using the empirical relationship found by Kieffer et al. (1973) for a standard value of the atmospheric pressure (4.5 Torr). Alternatively, following a more general approach, d can be determined for a generic value of the atmospheric pressure P using the equation found by Presley and Christensen (1997): k = (1.5 × 10−3 P 0.6 )d

(−0.11 log(

P 8.1×104

))

,

Table 2 Average grain sizes of the sediments in the sample basins, derived with the two methods P–C (Presley and Christensen, 1997) and Kieffer (Kieffer et al., 1973) described in the text. The corresponding water volumes, obtained for the two adopted values of the hydraulic radius, are also reported. Site No.

P–C diameter (mm)

Kieffer diameter (mm)

Water volume P–C H =1 m (km3 )

Water volume P–C H =5 m (km3 )

Water volume Kieffer H =1 m (km3 )

Water volume Kieffer H =5 m (km3 )

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20

0 .1 0 .3 8.8 0. 4 0.003 0 .4 0.003 0.04 0 .6 2.3 0. 3 1.5 1.5 2.0 1.0 0 .3 4.3 4.0 0. 4 0.3

0.2 0.2 2.0 0.3 0.02 0.3 0.02 0.06 0.3 0.7 0.2 0.5 0.5 0.7 0.4 0.2 1.1 1.0 0.3 0.2

43.1 2400 – 116 – 1440 – 22.1 5560 2380 225 9590 315 3130 235 3700 – – 468 59.5

15.0 852 – 42.5 – 523 – 7.6 2200 1410 62.8 4240 263 1500 102 1330 – – 178 20.9

28.7 1650 7820 64.0 3.5 795 63.0 49.6 2440 1400 115 8510 518 1370 106 2270 3970 920 232 30.6

10.0 581 4360 22.8 1.5 282 21.3 17.0 884 578 39.9 1920 96.0 540 39. 6 800 1840 428 83.6 10.6

(3)

This last quantity is linked to the height y of the site above the martian datum, by the relation:

Table 3 Relative abundances (by volume) of carbonates with respect to the other sediments present in the sample basins evaluated with the two methods P–C (Presley and Christensen, 1997) and Kieffer (Kieffer et al., 1973).

P = P 0 e − y /τ ,

Site No.

Carbonate abundance P–C (%) H =1 m

Carbonate abundance P–C (%) H =5 m

Carbonate abundance Kieffer (%) H =1 m

Carbonate abundance Kieffer (%) H =5 m

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20

0 .2 0.6 – 1 .4 – 1 .0 – <0.1 3 .3 5 .7 0.2 1 .2 1 .8 1 .1 2 .8 1 .0 – – 2 .8 0.2

0.1 0.2 – 0.5 – 0.3 – <0.1 1.3 3.4 0.1 0.6 0.9 0.5 1.2 0.4 – – 1.1 0.1

0 .3 0 .4 4 .0 0 .8 < 0.1 0 .5 0. 1 < 0.1 1 .4 3 .3 0 .1 0 .5 1 .0 0 .5 1 .3 0 .6 1 .0 1 .9 1 .4 0 .1

0 .1 0 .2 2.3 0 .3 <0.1 0 .2 <0.1 <0.1 0 .5 1.4 0 .1 0 .2 0 .3 0 .2 0 .5 0 .2 0.4 0 .9 0 .5 <0.1

where d is in micrometers, k is in W m−1 K−1 , and P is in Torr.

(4)

where P 0 is the atmospheric pressure at the martian datum (4.6 Torr) and τ is the scale height, which is 10.8 km for 210 K (Zurek et al., 1992). We remark that the relation between particle size and thermal conductivity derived by Presley and Christensen (1997) is significantly different from that estimated by Kieffer et al. (1973). For grain sizes in the narrow range 0.1–0.3 mm, similar values of the thermal conductivity are obtained with the two methods for the same pressure; however, relevant discrepancies are found outside this range of grain sizes. For each site, we first evaluate the thermal conductivity by means of Eq. (2) and TES thermal inertia measurements of the martian regolith, using the maps available at the web site of the University of Colorado (http://lasp.colorado.edu/marsdata/). Then, from the thermal conductivities we obtain two values, reported in Table 2, for the mean grain diameter: the first (hereinafter Kieffer diameter) obtained from the empirical relation derived by Kieffer et al. (1973) and the second (hereinafter P–C diameter) deduced from the analytical law (Eq. (3)), reported by Presley and Christensen (1997), coupled with Eq. (4). Between the two evaluations of the grain diameter, that of Presley and Christensen (1997) may be more reliable, since it takes into consideration the atmospheric pressure at each site under study. Unfortunately, while the Kieffer diameters derived for all the selected basins fall within the range that is suitable for Bagnold’s model (0.01–3.0 mm), the P–C diameters for five basins are outside this range. So, for these five basins we obtain the water volumes (see Table 2) and the carbonate percentages (Table 3) only using the Kieffer diameters. We note that the water volumes in Table 2 are not the real volumes of water present in the basins at any given time; they are estimates of the total volumes of all the water cycled through the basins, calculated as if it were all present at the same time.

On Mars as on Earth, atmospheric carbon dioxide dissolves in water to form carbonate and bicarbonate ions. In standing bodies of water, these ions tend to precipitate as sedimentary deposits of calcium and magnesium carbonates (e.g., CaCO3 and MgCO3 ). While many chemical and physical factors may limit the formation of carbonate sediments, an upper limit on the mass of carbonates that can be formed in a given volume of water can be derived from the number of Ca2+ and Mg2+ cations that are in solution. The concentration of cations in the early martian streams is not known. The Ca2+ and Mg2+ concentrations in terrestrial groundwaters in basaltic terrains range from 13 to 68 ppm and from 9 to 64 ppm, respectively (White et al., 1963). For a variety of other igneous rocks, the combined Ca2+ and Mg2+ concentrations rarely

Carbonate abundance in putative martian paleolakes

exceed 100 ppm (White et al., 1963). As in Goldspiel and Squyres (1991), we therefore use concentrations of 50 ppm for both calcium and magnesium cations in order to estimate the potential mass of carbonate precipitates that could have formed in the basins. It is worthwhile to note that the 100 ppm cation concentration assumed in the present work is consistent with the combined Fe2+ , Mg2+ , Ca2+ , Na+ , and K+ concentrations in Catling’s (1999) model of carbonate precipitation in martian lakes. Starting from these concentrations, the water volumes found with Bagnold’s model for each sample basin give (depending on the two values of the hydraulic radius H ) the carbonate abundances reported in Table 3. This table shows that the carbonate abundances in the chosen craters are about a few percent of the total sediment volume. Note that, due to the similar bulk density of carbonates and the other sediments (assumed equal to 2500 kg m−3 for both materials), the percentages in mass are practically coincident with those in volume. We recall that our approach for the computation of the carbonate abundances includes estimates of cation concentrations that are near the upper limits of what is observed in terrestrial groundwaters, and so may tend to overestimate the mass of the Ca2+ and Mg2+ cations supplied to the basins by the given volumes of inflowing water. This potential overestimation of cation mass at least partially balances the use of what are likely lower limits for the water volumes, leading, on the whole, to a reasonable estimation of the carbonate abundances. As discussed above, the carbonate abundance reported in Table 3 have been calculated assuming that the average size of the regolith on the floor of the craters, derived by thermal inertia measurements, is the same as that of the sediments deposited by the inlet channels during the past periods of fluvio-lacustrine activity. Actually, the particle sizes derived from thermal inertia values include the effects of aeolian and cratering processes that eventually occurred after the deposition of sediments into the basin, altering the original sediment size distribution. In order to take into account these effects, we consider also a more empirical approach that assumes for the sediments of all the basins an effective particle size of 0.7 mm, which is the average grain diameter measured by the Mars Exploration Rover Microscopic Imager at the landing site of Opportunity rover (Squyres et al., 2004b). The adopted value is also consistent with the grain sizes measured by the same kind of instrument at the landing site of Spirit rover (Herkenhoff et al., 2004). In this way we obtain the carbonate abundances reported in the last two columns of Table 4. As it can be seen, in the case of d = 0.7 mm, the abundances are generally higher than those obtained with diameters derived using P–C and Kieffer methods. This is due to the fact that the diameters obtained with the latter methods are generally smaller than 0.7 mm, perhaps due to aeolian resurfacing of the basin floor. This is particularly evident in the case of site No. 7 (lat. 0.0; long. 331.0◦ W) which shows a very high abundance of carbonates (about 12 or 5%, depending on the hydraulic radius) for the case of D = 0.7 mm. In spite of these large values, we do not think, however, that this site may be considered a good target in the search for carbonate deposits, since the small grain sizes derived from thermal inertia measurements (0.003–0.02 mm) strongly suggest the presence of an aeolian cover of the basin floor. The carbonate percentages reported in Tables 3 and 4 are in good agreement (except for the particular case of site No. 7—see above discussion) with the values (less than 6%) found by Goldspiel and Squyres (1991) for two basins in Mare Tyrrhenum SW and for other 32 sites selected in the ancient cratered terrains between the equator and 30◦ S with lateral sizes from ∼10 to ∼200 km. The agreement slightly worsens if we consider Gusev and a large basin in Margaritifer Sinus SE for which Goldspiel and Squyres (1991) calculated a larger carbonate percentage, between 5 and 9%, due to the fact that they used lower values for the average channel

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Table 4 Water volumes and relative abundances (by volume) of carbonates, with respect to the other sediments present in the sample basins, evaluated for an average grain size d = 0.7 mm and for the two adopted values of the hydraulic radius (see text). Site No.

Water volume (km3 ) H =1 m

Water volume (km3 ) H =5 m

Carbonate abundance (%) H =1 m

Carbonate abundance (%) H =5 m

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20

374 8724 4239 217 2162 2686 24,516 2521 6783 1398 638 5560 314 1373 187 11,661 2930 755 761 168

151 3516 1698 88 851 1073 10,707 1017 2806 572 249 2201 129 541 75 4759 1144 306 316 66

2 .2 2 .3 1 .9 2 .6 1 .1 1 .8 11.7 2 .2 4 .0 3 .3 0.8 0 .7 1 .1 0 .5 2 .2 3 .1 0.8 1 .6 4 .5 0 .5

0 .9 0.9 0.9 1.0 0 .4 0.7 5.1 0.9 1.6 1.4 0 .3 0 .3 0 .5 0.2 0 .9 1.3 0 .3 0 .6 1.9 0 .2

slopes, resulting in higher water volumes required to transport the same amount of sediments. As already found by Goldspiel and Squyres (1991), the chosen basins cannot account for the sequestration of significant quantities of CO2 from an early dense martian atmosphere. 5. Carbonate detection In order to evaluate the detectability of carbonates in the chosen sites via remote-sensing spectroscopy, the average percentages of these minerals in the basins (reported in Tables 3 and 4) have to be compared with the detection limits of the various spectrometers flown to Mars so far. Unfortunately, this quantity cannot be defined in a completely unambiguous way. The detection limit of a given particulate material mixed with other minerals depends, in fact, not only on the instrument, but also on surface texture and grain size, which both affect the strength of spectral features, as well as the geometrical and optical properties of the other components (Salisbury et al., 1991; Bishop et al., 1996; Wagner and Schade, 1996). Also the surface temperature plays an important role in this context, since an increase of the temperature facilitates the spectral detection of the material responsible for the feature (Fonti et al., 2001). Fine-grained carbonates, such as calcite, do have in the infrared (IR) region strong features that, in principle, are detectable for concentrations of a few percents (Wagner and Schade, 1996), or even less than 1% in ideal laboratory conditions (Bishop et al., 1996; Bandfield et al., 2003). However, this is true only if carbonates are not mixed with materials which can potentially mask the carbonate bands. For example some clay minerals exhibit bands near 2.35 μm (Bishop et al., 2008) that could overlap the carbonate feature at 2.3 μm, even if other carbonate bands, in particular that at about 3.9 μm, are not affected by this problem. Furthermore, field and laboratory works by Kirkland et al. (2000a, 2000b) have shown that the material from a deposit in Mormon Mesa, Nevada, with a carbonate concentration between 50 and 100%, has very low spectral contrast, probably due the surface texture of the sample which gives rise to cavity effects and volume scattering (Kirkland et al., 2000b). This prevents the identification of such deposit by means of airborne spectral measurements. According to Kirkland et al. (2000a, 2000b), this obser-

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vation raises the possibility that massive carbonate deposits similar to that in Mormon Mesa could also be present on Mars but escaped detection by the spectrometers flown to Mars so far, because of their weak bands. As far as such spectrometers are concerned, we note that CRISM (Compact Reconnaissance Imaging Spectrometer for Mars), onboard the Mars Reconnaissance Orbiter, and OMEGA, onboard Mars Express, have similar characteristics in the near IR, leading to comparable detection limits, of about 4% (Bibring et al., 2006). On the other hand for the TES spectrometer Christensen et al. (2001) and Stockstill et al. (2005) report a carbonate detection limit of 10–12% for coarse grains, while, according to Bandfield et al. (2003), the instrument has detected fine carbonate particles with an abundance as low as 2–5 wt%. This is an indication that the grain size may be important in this context, as discussed above. In light of the previous discussion, one can reasonably conclude that carbonates present in exposed aqueous deposits with concentrations less than a few percents could easily escape detection by the spectrometers flown to Mars so far. In the case of underground deposits, the situation is obviously much worse. Even if for all the chosen sites there is no evidence for volcanic resurfacing (Cabrol and Grin, 1999), there is the real possibility of aeolian coverage of the basin floor. In these cases, if the aeolian deposit is no more than a few centimeters deep, then a hypothetical underlying carbonate deposit should be detectable, not by remote-sensing spectral analyses, but by in situ measurements like those that will be performed during the upcoming NASA’s Mars Science Laboratory (MSL) mission. MSL, in fact, by means of a robotic arm will be able to acquire samples up to 5 cm deep in order to study them for example with the CheMin instrument, performing mineralogical analyses with the X-ray diffraction and X-ray fluorescence (XRD/XRF) standard techniques (Bish et al., 2007). 6. The case of Eberswalde crater An additional site considered in this study, not previously included in our original selection (Orofino et al., 2004), is the socalled Eberswalde crater, an elongated impact structure centered at about 24.2◦ S, 33.5◦ W (just north of Holden crater). This crater shows a well developed drainage basin, with 10 or more valleys contributing to a trunk valley that is about 90 km long and enters the crater through its northwest margin (see Fig. 1 of Malin and Edgett, 2003). Here the channel forms a 150 m thick complex deposit interpreted by Malin and Edgett (2003) as the exhumed and/or eroded remnant of a fluvial distributary fan. The Eberswalde delta comprises three lobes that form a body of sediment that measures roughly 10 km × 25 km, representing a volume of about 30 km3 (Bhattacharya et al., 2005). Since no spillways are observed, Eberswalde crater appears to be a closed basin for which our model is applicable. We include this crater, although very degraded, since it is one of the best studied (Malin and Edgett, 2003; Moore et al., 2003; Jerolmack et al., 2004; Bhattacharya et al., 2005; Lewis and Aharonson, 2006; Pondrelli et al., 2006). It is therefore possible to compare our calculation of the duration of the fluvio-lacustrine activity with the estimates given by other authors. For Eberswalde crater, due to the insufficient resolution of the currently available images, it is not possible to evaluate the sediment volume in the way described in Section 3. This volume, however, can be roughly estimated assuming a width of 1200 m and a depth of 40 m (Moore et al., 2003) for the 90-km-long main trunk, as well as a length of 20 km, a width of 500 m and a depth of 60 m (Moore et al., 2003) for each of the 7 main tributaries. In this way we obtain a sediment volume of about 9.3 km3 . As already noted by Moore et al. (2003), such an eroded volume is

much less than the volume of the fan (about 30 km3 in the updated evaluation by Bhattacharya et al., 2005). This means that the present morphology of the valleys is not representative of the past structure of the fluvio-lacustrine system, in the sense that this valley network, during the period of activity, was probably more developed and more sharply incised in the ground with respect to what is shown by the present-day topography. Consequently in our calculations of the water volume and of the carbonate percentage, we conservatively use a value of 30 km3 for the sediment volume. In other words, we assume that all the sediments carried by the inflow valley are concentrated in the fan. As in the case of the other craters, in order to perform our calculations, it is necessary to evaluate the hydraulic radius of the inlet channel, which today is not visible but presumably was once present inside the main valley of the fluvial network. Considering the relatively small width of this valley (1200 m, according to Moore et al., 2003), a correspondingly small width of the interior channel seems reasonable; for such width we can assume 100 m  w  400 m, i.e. its value ranges between the minimum and the average widths of the 21 interior channels studied by Irwin et al. (2005). Using the above quoted relation H = ξ w 0.69 (Williams, 1988), the corresponding hydraulic radius is in the range 1–3 m. For these two values, and for a slope of 0.015 (obtained from MOLA data) and a grain size of 2.9 mm (derived from the mean thermal inertia of the crater, using the P–C relationship between thermal inertia and grain size), our model indicates that 8590 and 6070 km3 of water, respectively, are necessary in order to transport the 30 km3 of sediments present in the fan. With these water volumes, following the same approach used for the other craters, we obtain a carbonate abundance of about 3.4 and 2.4%, respectively. On the other hand, if we use a sediment grain diameter of 0.7 mm, we obtain a carbonate percentage of about 1.2 and 0.7% for the two hydraulic radii of 1 m and 3 m. These values are similar to those found for the other sites (reported in Tables 3 and 4), so that the same conclusions drawn in the previous section about the carbonate abundances in the sample craters also hold for Eberswalde basin. At this point two methods can be envisioned to estimate the duration of the fluvio-lacustrine activity in the Eberswalde basin. 6.1. Method 1 For the above reported input parameter values and for a grain diameter of 2.9 mm, our model gives a sediment mass j transported per unit channel width and per unit time in the range 27–240 kg m−1 s−1 . The upper limit is obtained for a 3 m deep (400 m wide) channel, while the lower limit is the result for a channel depth of 1 m (width of 100 m). This implies that the sediment mass J = j · w transported per unit time is in the range 2.7 × 103 –9.4 × 104 kg s−1 , depending on the assumed width of the channel. By comparison, the sediment transport capacity of Huang He river, one of the most efficient terrestrial rivers in this respect, would be 9.0 × 104 kg s−1 when scaled to the martian gravity. Given a sediment mass M S = 7.5 × 1013 kg (obtained from the fan volume with an assumed bulk density of 2500 kg m−3 ), we find that the sediment transport would require a time t 1 = M S / J = 25–880 yr. It is necessary to remember, however, that this evaluation assumes that no work is being done by the flow in eroding the source material or breaking the sediment up to small sizes necessary for the transport, so that the above reported values have to be regarded as a conservative estimate for the time required to carve the valley systems and, then, a lower limit for the duration of the fluvio-lacustrine activity in Eberswalde crater.

Carbonate abundance in putative martian paleolakes

6.2. Method 2 Alternatively one can follow the approach by Lewis and Aharonson (2006), which is based on the evaluation of the water volume and an assumed value for the loss rate of water from the basin (L = dh/dt), mainly due to evaporation and/or percolation into the ground. Starting from the estimated mean area of the paleolake in the crater ( A = 240 km2 , according to Lewis and Aharonson, 2006), the water volume (V w = 6070–8590 km3 ) given by our model for d = 2.9 mm implies an effective water depth h = V w / A of the order of 3 × 104 m. As discussed before, this is the total depth of all the water that was cycled through the basin, calculated as if it were all present at the same time. In any case, using this water depth it is possible to evaluate the lifetime t 2 of the lake, if the loss rate of water from the basin L is known. Following Irwin et al. (2005) and Lewis and Aharonson (2006), we can assume for comparison a value L = 1–10 m/yr, loss rates typical of terrestrial lakes, so that t 2 = h/ L ≈ 3 × 103 –3 × 104 yr. Our results can be compared with previous estimates of the duration of the fluvial activity done by other authors and based on the evaluation of the formation time for the fan: Moore et al. (2003) find this time to be in the range 2 × 103 –1 × 106 yr, for assumed values of sediment yield typical of terrestrial drainage basins; Jerolmack et al. (2004) estimate a minimum formation time for the delta of 50 yr; Bhattacharya et al. (2005) suggest that the lake in Eberswalde crater persisted for at least 1.5 × 105 yr, while Lewis and Aharonson (2006) find a minimum formation time for the delta of several hundred years. We note that our estimate of t 1 is in good agreement with the values obtained by Jerolmack et al. (2004) and by Lewis and Aharonson (2006), while our evaluation of t 2 is within the range found by Moore et al. (2003) but less than the value suggested by Bhattacharya et al. (2005). We also note that discrepancies between our values of t 2 and those obtained by Lewis and Aharonson (2006) are only due to the fact that these authors assume a sediment volume of 6 km3 (Malin and Edgett, 2003) and a water/sediment volume ratio of 30, while we assume a sediment volume of 30 km3 (see above) and use water/sediment volume ratios of 200 and 290 obtained from our model for H = 3 m and H = 1 m, respectively. Repeating the calculations for a grain size d = 0.7 mm, we obtain t 1 = 7–300 yr and t 2 = 7 × 102 –1 × 104 yr, with a poorer agreement with the literature. In any case, we can conclude that the fluvio-lacustrine activity in Eberswalde would have lasted at least for a period on the order of a thousand years for a continuous flow, or one or more orders of magnitude longer for more intermittent flows. This last possibility is quite likely, especially if we take into consideration the cases of other paleo-fluvial martian systems, such as Ma’adim Vallis (Cabrol et al., 1996, 1998), Warrego Valles (Ansan and Mangold, 2006), and the above quoted channel in Libya Montes, studied by Jaumann et al. (2005). All of these systems should have experienced various episodes of water flow interrupted by long periods of inactivity. A possible explanation for the episodic flow is that cyclic variations in the obliquity and the orbit eccentricity could have produced large climate changes; the latter, in turn, could have altered the hydrological cycles of these fluvial systems, switching them on and off alternatively (Cabrol et al., 1998). 7. Discussion and conclusions In this work we apply the Bagnold (1966) sediment transport model to a sample of putative martian paleolacustrine basins in order to estimate the volume of water necessary to remove the sediments and carve the inflow valleys. From this quantity we evaluate, on the basis of reasonable assumptions for cation con-

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centrations in martian groundwaters, the abundance of carbonates that could have precipitated in these basins. We remark that for the hydraulic radii (1–5 m), channel slopes (0.006–0.05) and sediment grain sizes (0.01–3.0 mm) considered here, the thickness of the bedload zone is negligible in comparison to the water height and the streams can be treated as fully turbulent. In other words the conditions at the basis of Bagnold’s model are met. We recall, however, that the topography of these basins is generally too complex to be well assessed with the current data; thus, estimating water and sediment volumes and the duration of standing water bodies is subject to some uncertainties. There are also many uncertainties in the sediment transport calculations. Kleinhans (2006) has reviewed several sediment transport models, including the Bagnold (1966) model used here, and found that the Bagnold’s model may significantly underestimate the water volumes needed to transport sediment. It is important to remember, however, that all water volume estimates derived from sediment transport calculations are really lower limits since these calculations only treat the transport of the sediment and do not include the additional water flow needed to erode the regolith and produce the sediment in the first place. As more detailed information is collected on sediment and bed forms at specific sedimentation sites, the alternative equations presented by Kleinhans (2006) may offer a good opportunity to better understand the flow regime for each location. For a general application such as this study, however, the standard Bagnold model provides a sufficient means to estimate the potential for carbonate mineral sedimentation at different sites on Mars. We address the uncertainties associated with sediment transport on Mars by using a range of stream depths in our calculations. It is important to note that the carbonate abundances presented here are calculated with the assumption that the carbonate precipitates and the fluvially transported clastic sediment are uniformly mixed on the basin floor. In an undisturbed stream-fed lake system, it is reasonable to expect a degree of vertical and horizontal heterogeneity of chemical and clastic sediments since incoming clastic sediments generally would not settle at the same rate chemical sediments precipitate, and following Stoke’s Law, larger clastic sediment grains would generally settle out faster and nearer to the inflow valley than finer clastic sediment grains. In cases of episodic water flows into a lake basin, a series of clastic-chemical sediment layers could be anticipated as each flow episode could result in alternating strata of chemical precipitates and clastic sediments. A uniformly mixed collection of fluvial sediments may therefore not be a good model for a pristine sub-aqueous sedimentary structure. However, in cases of basins that have been physically disturbed, such as those subjected to impact gardening (Hartmann et al., 2001), the assumption of uniform sediment mixing is a reasonable test case to compare with the limiting cases of fully exposed carbonate sediment layers (which have not been detected from orbit), fully buried carbonate sediment layers (which cannot be directly detected from orbit), and no carbonate sediments at all. The uniform mixing results from this study are consistent with the conclusion of Hartmann et al. (2001) that, unless preserved through special conditions, ancient carbonate deposits would be diluted through mixing with other sediments to a point where such deposits could not be detected by the instruments aboard MGS and Mars Express. The results of our computations for Eberswalde crater and for the other sample craters (see Tables 3 and 4), indicate that, except for the already discussed site No. 7, the expected carbonate abundances in these basins are less than a few percent, that is less than (or at best comparable with) the detection limit of the spectrom-

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eters TES, OMEGA and CRISM. These results could explain why all these instruments failed to detect carbonate deposits on Mars. Since the chosen basins are among the most likely sites of carbonate precipitation on Mars, this negative result makes it unlikely that carbonate deposits could be detected elsewhere on the martian surface by the current generation of orbiting spectrometers. Our results also indicate that the lack of positive detection of carbonate deposits on Mars cannot be used as an argument against the presence of a sustained ancient hydrological cycle on the planet. In fact, as already found by Goldspiel and Squyres (1991), the quantity of water required to transport the sediments in the analyzed basins is greater than that which could be produced by a single discharge of the associated aquifer, unless the material of the martian soil was very fine-grained and non-cohesive to a depth of hundreds of meters. This means that aquifer recharge is probably required to erode the inflow channels, so that a hydrological cycle was likely present on Mars when the fluvio-lacustrine systems formed. Such a cycle, able to move large volumes of water and to create lakes, could have been active intermittently on Mars in the past, producing carbonate deposits that could escape detection by the instruments that have flown to date. Acknowledgments We thank Janice Bishop and Joshua Bandfield for their comments and remarks which helped us to improve the quality of the paper. This research has been partially supported by Italian Space Agency (ASI) and the Italian Ministry of University and Research (MUR). References Ansan, V., Mangold, N., 2006. New observations of Warrego Valles, Mars: Evidence for precipitation and surface runoff. Planet. Space Sci. 54, 219–242. Bagnold, R.A., 1966. An approach to the sediment transport problem from general physics. US Geol. Survey Prof. Paper 422-I. Baker, V.R., Strom, R.G., Gulick, V.C., Kargel, J.S., Komatsu, G., Kale, V.S., 1991. Ancient oceans, ice sheets and the hydrological cycle on Mars. Nature 352, 589–594. Bandfield, J.L., Glotch, T.D., Christensen, P.R., 2003. Spectroscopic identification of carbonate minerals in the martian dust. Science 301, 1084–1087. Bhattacharya, J.P., Payenberg, T.H.D., Lang, S.C., Bourke, M., 2005. Dynamic river channels suggest a long-lived Noachian crater lake on Mars. Geophys. Res. Lett. 32, doi:10.1029/2005GL022747. L10201. Bibring, J.-P., Langevin, Y., Gendrin, A., Gondet, B., Poulet, F., Berthé, M., Soufflot, A., Arvidson, R., Mangold, N., Mustard, J., Drossart, P., and the OMEGA Team, 2005. Mars surface diversity as revealed by the OMEGA/Mars Express observations. Science 307, 1576–1581. Bibring, J.-P., Langevin, Y., Mustard, J.F., Poulet, F., Arvidson, R., Gendrin, A., Gondet, B., Mangold, N., Pinet, P., Forget, F., 2006. Global mineralogical and aqueous Mars history derived from OMEGA/Mars Express data. Science 312, 400–404. Bish, D.L., Blake, D., Sarrazin, P., Treiman, A., Hoehler, T., Hausrath, E.M., Midtkandal, I., Steele, A., 2007. Field XRD/XRF mineral analysis by the MSL CheMin instrument. Lunar Planet. Sci. XXXVIII. Abstract 1163. Bishop, J.L., Koeberl, C., Kralik, C., Fröschl, H., Enolert, P.A.J., Andersen, D.W., Pieters, C.M., Wharton, R.A., 1996. Reflectance spectroscopy and geochemical analyses of Lake Hoare sediments, Antarctica: Implications for remote sensing of the Earth and Mars. Geochim. Cosmochim. Acta 60, 765–785. Bishop, J.L., Lane, M.D., Dyar, M.D., Brown, A.J., 2008. Reflectance and emission spectroscopy study of four groups of phyllosilicates: Smectites, kaolinite-serpentines, chlorites and micas. Clay Miner. 43, 35–54. Bullock, M.A., Moore, J.M., 2007. Atmospheric conditions on early Mars and the missing layered carbonates. Geophys. Res. Lett. 34, doi:10.1029/2007GL030688. L19201. Cabrol, N.A., Grin, E.A., 1999. Distribution, classification, and ages of martian impact crater lakes. Icarus 142, 160–172. Cabrol, N.A., Grin, E.A., Dawidowicz, G., 1996. Ma’adim Vallis revisited through new topographic data: Evidence for an ancient intravalley lake. Icarus 123, 269–283. Cabrol, N.A., Grin, E.A., Landheim, R., 1998. Ma’adim Vallis evolution: Geometry and models of discharge rate. Icarus 132, 362–377. Carr, M.H., 1981. The Surface of Mars. Yale Univ. Press, New Haven. Carr, M.H., 1996. Water on Mars. Oxford Univ. Press, New York. Catling, D.C., 1999. A chemical model for evaporites on early Mars: Possible sedimentary tracers of the early climate and implications for exploration. J. Geophys. Res. 104, 16453–16470.

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