Chemical models for martian weathering profiles: Insights into formation of layered. phyllosilicate and sulfate deposits. Mikhail Yu.

Transcription

1 Chemical models for martian weathering profiles: Insights into formation of layered phyllosilicate and sulfate deposits Mikhail Yu. Zolotov* School of Earth and Space Exploration, Arizona State University, Tempe, AZ , USA. Phone: (480) , Fax: (480) , Mikhail V. Mironenko Vernadsky Institute of Geochemistry and Analytical Chemistry, Russian Academy of Sciences, 19 Kosygin Str., Moscow , Russia. Submitted to Icarus on March 5, 2015 Accepted for publication on April 11, 2016 DOI information: /j.icarus *Corresponding author Highlights: Weathering models reproduce clay-bearing Noachian stratigraphies Acid sulfate weathering is more likely than weathering at a dense CO 2 atmosphere Sulfates and chlorides are by-products of Noachian acid weathering Some Hesperian sulfate deposits formed through remobilization of Noachian salts 1

2 Running title: Martian weathering profiles Abstract. Numerical chemical models for water-basalt interaction have been used to constrain the formation of stratified mineralogical sequences of Noachian clay-bearing rocks exposed in the Mawrth Vallis region and in other places on cratered martian highlands. The numerical approaches are based on calculations of water-rock type chemical equilibria and models which include rates of mineral dissolution. Results show that the observed clay-bearing sequences could have formed through downward percolation and neutralization of acidic H 2 SO 4 -HCl solutions. A formation of weathering profiles by slightly acidic fluids equilibrated with current atmospheric CO 2 requires large volumes of water and is inconsistent with observations. Weathering by solutions equilibrated with putative dense CO 2 atmospheres leads to consumption of CO 2 to abundant carbonates which are not observed in clay stratigraphies. Weathering by H 2 SO 4 -HCl solutions leads to formation of amorphous silica, Al-rich clays, ferric oxides/oxyhydroxides, and minor titanium oxide and alunite at the top of weathering profiles. Mg-Fe phyllosilicates, Ca sulfates, zeolites, and minor carbonates precipitate from neutral and alkaline solutions at depth. Acidic weathering causes leaching of Na, Mg, and Ca, from upper layers and accumulation of Mg-Na-Ca sulfate-chloride solutions at depth. Neutral MgSO 4 type solutions dominate in middle parts of weathering profiles and could occur in deeper layers owing to incomplete alteration of Ca minerals and a limited trapping of Ca to sulfates. Although salts are not abundant in the Noachian geological formations, the results suggest the formation of Noachian salty solutions and their accumulation at depth. A partial freezing and migration of alteration solutions could have separated sulfate-rich compositions from low-temperature chloride brines and contributed to the observed diversity of salt deposits. A Hesperian remobilization and release of subsurface MgSO 4 type solutions into newly-formed depressions 2

3 could account for formation of massive layered sulfate deposits through freezing or evaporation. This scenario explains the observed deficiency of salts in Noachian formations, a paucity of Hesperian phyllosilicates, and the occurrence of sulfate deposits in Valles Marineris troughs, chaotic terrains, and some craters of the Hesperian age. Keywords: Mars, surface; Mineralogy 1. Introduction The occurrence and composition of minerals deposited from aqueous solutions (salts, phyllosilicates, zeolites, opaline silica, etc.) could be used to constrain paleoenvironments and history of exogenic processes on Mars. Orbital spectral observations of the martian surface indicate uneven occurrences of aqueously-formed minerals in geological formations of different ages. Secondary minerals seen in the Noachian (3.7 to 4.1 Ga) terrains are dominated by phyllosilicates, while the Hesperian (3.0 to 3.7 Ga) formations are characterized by massive deposits of layered sulfates observed in the Valles Marineris trough system, chaotic terrains, at the Meridiani Planum, and several other places (Poulet et al., 2005; Gendrin et al., 2005; Mustard et al., 2008; Murchie et al., 2009a; Carter et al., 2013). The initial interpretation of these data suggested deposition of phyllosilicates from Noachian alkaline aqueous solutions and formation of sulfates from acidic fluids related to the Hesperian volcanism (Bibring et al., 2006). Subsequent detections of coexisting sulfates and phyllosilicates in several locations imply more complex formation scenarios (Wray et al., 2009, 2011; Milliken et al., 2010; Noe Dobrea et al., 2012). Milliken et al. (2009) noticed a paucity of Noachian salts and/or oxides which were expected by-products of the phyllosilicate formation. In this paper, we note that a formation of 3

4 massive sulfate deposits through acidic alteration of mafic rocks in the Hesperian epoch (Bibring et al., 2006) would have led to corresponding volumes of acid-resistant Si-Al-rich minerals, which are not seen in formations of that age. We argue for the formation of sulfate solutions together with layered Noachian phyllosilicates followed by aqueous remobilization and deposition of sulfates in the Hesperian epoch. Fe-Mg phyllosilicates (smectites, chlorites) are the most abundant clay minerals detected on Mars (e.g. Ehlmann et al. 2011; Carter et al., 2015; Sun and Milliken, 2015). A significant fraction of clay minerals occurs in layered geological formations and is denoted as phyllosilicates in stratigraphies by Ehlmann et al. (2011). In about half of the occurrences (Carter et al., 2015), Fe-Mg phyllosilicates in stratigraphies are overlain by a layer with Al-rich clays such as smectites (montmorillonite, beidellite) and kaolinite-group minerals (kaolinite, halloysite). With some notable exceptions, an average thickness of layered clays is about several meters (Carter et al., 2015). The typical thickness of Al-rich layers (from a meter to a few tens of meters) is less than the thickness of underlying Fe-Mg clay units by a factor of 2 10 (e.g. Bishop et al., 2008; Le Deit et al., 2012; Carter et al., 2015). This stratigraphic sequence of clay minerals has been observed in ~120 locations of the martian highlands in which Noachian bedrocks are exposed due to erosion, tectonics, or impact cratering (Carter et al., 2015). The occurrence of stratified clay-rich formations do not show a correlation with a geological unit related to an endogenic process. As examples, the clay sequence is reported in the Arabia Terra/Mawrth Vallis region (Noe Dobrea et al., 2010), in Noachis Terra (Wray et al., 2009), in Nili Fossae (Ehlmann et al., 2009; Gaudin et al., 2011), in western Arabia Terra (Noe Dobrea et al., 2010), in Eridania basin (Noe Dobrea et al., 2010), in South Terra Sabaea (Carter et al., 2015), and in the 4

5 Valles Marineris and Thaumasia Planum regions where plateau rocks are exposed in chasmata, pits, valleys, and crater walls (Le Deit et al., 2012; Murchie et al., 2009a). The occurrence of similar stratigraphic sequences of clay minerals in the equatorial to midlatitude highlands with similar ages indicates a widespread or even global rock alteration occurred ~ Ga in the middle to late Noachian epoch (Carter et al., 2015). The similarity of observed clay stratigraphies with those in weathering profiles on terrestrial basalts is consistent with in situ alteration of rocks by aqueous solutions (pedogenesis) (Bibring et al., 2006; Bishop et al., 2008; Ehlmann et al., 2009; Noe Dobrea et al., 2010; McKeown et al., 2009; Loizeau et al., 2007, 2010; Gaudin et al., 2011; Le Deit et al., 2012; Greenberger et al., 2012; Carter et al., 2013, 2015). The pedogenesis implies percolation of low-temperature surface waters to the depth and leaching cations (Mg, Fe, Ca, Na, etc.) from near-surface rock layers. Early-formed secondary phases (e.g. Mg-Fe phyllosilicates) are gradually replaced by lowsolubility Al-clays. A weathering formation scenario for martian clay-bearing stratigraphies is consistent with gradual transitions between upper Al-rich layers, Mg-Fe-rich clay units, and partially altered/unaltered rocks, with often occurrences of the stratigraphies on highlands, as well as with the lack of detection of phases that form at elevated temperatures (Carter et al., 2015). The weathering scenario is consistent with a usual presence of primary minerals from basalts (pyroxenes, plagioclases) in clay-bearing mineral assemblages (Poulet et al., 2008b). Geochemical consequences of the supposed Noachian weathering remain to be evaluated in the contexts of origin, composition, and evolution of alteration fluids, and a fate of leached elements. A validity of the weathering hypothesis could be better tested through in situ investigation of weathering profiles with further rovers. At present, insights could be obtained from a comparison of martian well-exposed clay-bearing stratigraphic sequences with terrestrial 5

6 weathering profiles (e.g. Gaudin et al., 2011; Greenberger et al., 2012) and with results of experimental and numerical modeling. A comparison could be done for the Arabia Terra/Mawrth Vallis region, which is the most remotely-studied area with phyllosilicates in stratigraphies. In the Mawrth Vallis region, an exceptionally thick (> m) sequence of secondary minerals is observed in flanks of the valley and other erosion features over an area of ~300 by ~400 km (Loizeau et al., 2007, 2010). The upper stratigraphic unit (10 50 m thick) contains vol.% of kaolinite-group minerals and/or montmorillonite-like Al-rich clays, and a lower unit contains vol.% of Mg-Fe smectites (Poulet et al., 2005, 2008b; Loizeau et al., 2007, 2010; Wray et al., 2008; Bishop et al., 2008, 2013; McKeown et al., 2009, 2011; Noe Dobrea et al., 2010; Michalski et al., 2013; Viviano and Moersch, 2013). Amorphous Si-Al hydrous oxide phases (hydrated silica, allophane; Bishop et al., 2008, 2013; Bishop and Rampe, 2014; Michalski et al., 2013), imogolite (Bishop and Rampe, 2014), patchy occurrences of jarosite and copiapite (Farrand et al., 2009, 2014; Bishop et al., 2013; Michalski et al., 2013), and Fe 3+ oxides/oxyhydroxides (hematite, ferrihydrite; Wray et al., 2008; McKeown et al., 2009; Farrand et al., 2009) were reported in the upper unit. A layer enriched in a ferrous phyllosilicate occurs locally below the Al-rich unit (Bishop et al., 2008, 2013; McKeown et al., 2009). Zeolites and/or sulfates could be present within the lower Fe-Mg clay unit (Michalski and Fergason, 2009; Bishop et al., 2013; Michalski et al., 2013). A Ca sulfate was reported in the lower part of that unit (Wray et al., 2010). Although Fe 3+ oxides/oxyhydroxides were reported in the region, their genetic relationship with other phases in stratigraphies is unclear (Poulet et al., 2008b; Michalski et al., 2013; Bishop et al., 2013). Plagioclases and pyroxenes are detected throughout the sequence, especially in deeper stratigraphic layers (Poulet et al., 2008b; Viviano and Moersch, 2013). These data imply incomplete alteration of basaltic material in the supposed alteration 6

7 sequence. Although several formation scenarios were proposed for the stratigraphic sequence in the region (Michalski et al., 2013), chemical weathering is considered as a reasonable pathway. In this paper, we used numerical physical-chemical approaches to get insights into chemical weathering of basaltic materials in the Noachian epoch. We constrained the development of weathering profiles, evaluated the composition and ph of formed aqueous solutions, assessed compositional changes during subsequent freezing of subsurface fluids, and proposed a coupled formation of aqueous minerals in observed clay- and sulfate-rich stratigraphies. 2. Numerical models for weathering profiles We have considered alteration of a permeable basaltic material by surface aqueous solutions which percolated to the depth. The models took into account a thermodynamicallycontrolled precipitation of secondary phases from aqueous solution and an equilibration among components in solution. The initial rock was represented by the composition of martian Adirondack-type olivine basalt (Table 1). Initial fluids were characterized by pure water, solutions equilibrated with current atmospheric CO 2 (5.3 mbar) and O 2 (7.6 μbar), and an acidic H 2 SO 4 -HCl solution, which could have formed through volcanic or impact-induced degassing. Some initial fluids were equilibrated with dense atmospheres ( bars CO 2 ) containing current or elevated partial pressures of O 2. The chosen acidic fluid was a 0.5 molal H 2 SO 4, 9.6 millimolar HCl (ph = 1.2) solution adopted from Zolotov and Mironenko (2007a). The solution corresponds to the S/Cl molar ratio of 5.2 that reflects the composition of martian soils at landing sites of Viking 1 and 2, Mars Pathfinder, and the Spirit rover. The modeling was performed for 0 o C at bar total pressure in the O-H-C-S-Cl-Mg-Fe-Ca-Al-Si-Na-K-Ti-Mn water-rock- 7

8 gas type closed system. Some runs for a surface rock layer were performed in the system open with respect to atmospheric CO 2 and O 2. Formation of methane was suppressed because of its slow interaction with inorganic carbon species at the chosen temperature (T) and pressure (P). Three modeling approaches have been applied (Table 2). In models I and II, weathering was quantified through calculations of complete chemical equilibria in the systems. The input data included the bulk chemical composition of the system and Gibbs free energies of formation at the selected T and P of chemical species which could form in the system. Activity coefficients of aqueous species were calculated with the extended Debye-Hückel model. The results provided data on the equilibrium speciation of solid, aqueous, and gaseous phases together with activities of solutes and the ph. Chemical equilibria were calculated through minimization of the Gibbs free energy of the system with GEOCHEQ codes (Mironenko et al., 2008). The codes, sources of thermodynamic data for species, and solid solution models were described elsewhere (Zolotov et al., 2006; Zolotov and Mironenko, 2007a). Montmorillonite and saponite were presented by ideal solutions of Mg-, Ca-, Na-, K-, and Fe-bearing end-members. The compositions of brines and precipitated salts formed through equilibrium or fractional freezing of solutions were calculated with the FREZCHEM code (Marion et al., 2010) which uses Pitzer model to calculate activities of water and solutes at elevated ionic strengths. In Model I, chemical equilibria were calculated in a series of water/rock (W/R) mass ratios. The results could be interpreted by different means. The first approach implies that the W/R ratio decreases with depth in a weathering profile and a decreasing W/R ratio is considered as a proxy for depth. Indeed, deeper layers are typically affected by lesser masses of solutions because of lowering porosity, permeability, and a consumption of water in upper horizons through hydration of the rock. In another approach, the calculations give insights into alteration progress within a 8

9 chosen rocky layer. An elevated W/R ratio corresponds to initial stages of alteration when a small mass of rock reacts with the whole mass of aqueous solution. The progress of alteration corresponds to a decreasing W/R ratio as more rock involves in reactions. In this interpretation, the calculated secondary mineralogy corresponds to the mass of reacted rock rather than to the whole rock in the layer. In Model II, we considered alteration of rock layers by a series of percolating solution parcels. The increasing number of water solution parcels corresponds to alteration progress. The uppermost rock layer interacted with solutions of a chosen initial composition. In a typical run, 1 kg of basalt in the upper layer sequentially reacted with 250 solution parcels (1 kg per parcel). Each other dimensionless rock layer sequentially reacted with solution parcels formed through water-rock interactions in upper layers. This model allowed us to assess changes in secondary mineralogy and solution composition within each layer, as well as a progression of the weathering profile to depth. In Model III, alteration of rock s layers by percolating fluids was modeled by taking into account dissolution rates of mineral grains. This approach is based on assumptions that formation of secondary minerals is preceded by dissolution of primary phases, mineral dissolution is the time-limiting step of rock alteration, and all species in solution are in chemical equilibrium with each other (Helgeson, 1968). The model assumes formation of solution composition through kinetically-controlled dissolution of unstable primary and previously formed secondary phases, and thermodynamically-controlled precipitation of secondary minerals. At each time step in a rock layer, dissolving primary and/or secondary phases coexist with secondary phases which are in equilibrium with aqueous solution. In each layer, solution is allowed to react with rocks during a specified period of time. In addition to data required to calculate equilibria, the input 9

10 information includes grain sizes and shapes of mineral grains, and a fraction of mineral surfaces exposed to solution. A static version of this kinetic-thermodynamic approach is described elsewhere (Zolotov and Mironenko, 2007a; Mironenko and Zolotov, 2012). The flow model has procedures to quantify ph- and T-dependent rates of mineral dissolution coupled with chemical equilibria in solution and percolation of fluids. We used dissolution kinetic constants composed in (Zolotov and Mironenko, 2007a). The used mineralogy of the initial rock (Table 1) had no sulfides, and a minor contribution of sulfides into the sulfur balance has not been considered in this model. 3. Results All calculations led to weathering profiles in which Al-rich phyllosilicates formed above less leached materials with abundant Fe-Mg clay minerals. In each model, the composition of major secondary minerals and appearances of minor alteration phases were somewhat different and depended on the type of initial fluid and degree of weathering. Weathering by initial pure (neutral) and slightly acidic fluids equilibrated with current atmospheric CO 2 did not produce amorphous silica and abundant sulfates. Weathering by these solutions required significantly larger masses of water to develop Al-rich phyllosilicates than alteration by more acidic fluids. Weathering by solutions equilibrated with dense CO 2 atmospheres led to abundant carbonates in middle parts of profiles which have not been reported in martian clay-bearing stratigraphies. Weathering by H 2 SO 4 -HCl type acidic solutions led to the formation of weathering profiles with abundant amorphous silica, Ca sulfates, and other phases reported in the Mawrth Vallis region. Therefore, for Models II and III, we reveal results obtained for H 2 SO 4 -HCl type solutions. 10

11 3.1. Weathering models at the variable W/R ratio Equilibrium calculations with initial H 2 SO 4 -HCl acidic solutions demonstrated a significant leaching of Na, Mg, Ca, and Fe from basalt at elevated W/R ratios which could represent upper layers of a weathering profile and an advanced stage of alteration in these layers (Figs. 1a and 1b). Potassium is leached only at high W/R ratios, and Al, Si, and Ti are the least mobile elements. In terms of depth, the modeled uppermost layer mostly consists of amorphous silica, kaolinite, and minor Ti oxide (rutile). An unabundant alunite is present in association with montmorillonite, kaolinite, and amorphous silica. With increasing depth, dominated secondary phases are montmorillonite and Fe-Mg phyllosilicates, in that order. A montmorillonite-rich assemblage contains amorphous silica, kaolinite, and minor gypsum. The observed formation of Fe-chlorite indicates the thermodynamic drive for its precipitation, though the mineral may not precipitate at 0 C by kinetic reasons (section 4.1). Montmorillonite and amorphous silica become less abundant at lower W/R ratios (with increasing depth) but they are present in assemblages dominated by Fe-Mg phyllosilicates. Zeolites (stilbite, stellerite, analcime) and gypsum are present among Fe-Mg phyllosilicates and coexist with alkaline solutions at low W/R ratios at depth (Figs. 1a and 1c). In terms of alteration progress, the modeled mineralogy suggests sequential alteration of earlier formed secondary phases (Fe-Mg clay minerals, zeolites, and gypsum) into Al-rich clay minerals and amorphous silica. The formation of Al-rich clays and amorphous silica at elevated W/R ratios could also be referred to secondary minerals which form at initial stages of acidic weathering when only a small portion of basalt is dissolved without a strong effect on the solution ph (cf. Zolotov and Mironenko, 2007a). For the chosen initial H 2 SO 4 -HCl type solution, basalt alters to the kaolinite-silica-rutile mineralogy at the W/R mass ratio of ~150. In contrast, 11

12 weathering by S-, Cl-free water at comparable W/R ratios ( ) leads to mineralogy rich in Fe-Mg phyllosilicates and zeolites (Fig. 1d). Only Na is leached and the solution ph is strongly alkaline (Figs. 1e and 1f). Significant amounts of solutions (W/R > 10 4 ) are needed to form alteration assemblages with ~50% of Al-rich phyllosilicates. Amorphous silica does not form even at W/R = 10 5 and kaolinite is replaced by gibbsite at higher W/R ratios. Alteration by S-, Cl-free solutions equilibrated with denser CO 2 atmospheres produces Alrich phyllosilicates at lower W/R ratios (Figs. 2 and 3). Kaolinite, montmorillonite, and amorphous silica form in basalts altered by CO 2 -rich mildly acidic solutions. Weathering by solutions equilibrated with 1 bar CO 2 leads to montmorillonite-rich assemblages at W/R mass ratio as low as 50. Siderite is a major mineral in these assemblages formed from partially neutralized fluids at ph = 6 7, which could characterize a shallow subsurface environment. Calcite forms at neutral and alkaline phs in middle parts of weathering profiles. Lower parts of the profiles at low W/R ratios and ph = 7 12 are presented by Fe-Mg phyllosilicates, zeolites, minor gypsum, and unaltered minerals from basalt. These results indicate that a dense CO2 atmosphere does not prevent formation of abundant Fe-Mg phyllosilicates, which cannot form at the planetary surface, however (cf. Chevrier et al., 2007). The presence of a minute amount of O 2 (~ mole per kg H 2 O expected at current surface solutions) in these models contributes to the oxidation of sulfides at elevated W/R ratios but does not affect reduced Fe and S mineralogy in the middle and lower parts of weathering profilers. Calculations performed for O 2 -saturated solutions at the current atmospheric 0.13 vol. % O 2 scaled to dense CO 2 atmospheres led to analogous reduced weathering profiles. Modeled weathering by H 2 SO 4 -HCl solutions saturated with dense CO 2 atmospheres led to formation of stratified clay mineralogy and abundant carbonates together with Ca sulfates in the middle parts of weathering profiles. 12

13 The calculated speciation of aqueous solutions in equilibrium with secondary minerals is different for weathering by H 2 SO 4 -HCl and S-, Cl-free solutions (Fig. 4). Acidic weathering leads to formation of low-ph sulfate-chloride solutions with elevated concentrations of Fe 2+ and Mg 2+ species in upper parts of weathering profiles (Fig. 4a), though solution-flow models indicate transfer of these cations to lower layers and their partial trapping in phyllosilicates (Section 3.2). As the W/R ratio decreases with depth, near-neutral MgSO 4 type solutions become dominant. As W/R decreases further, Mg sulfate solutions are replaced by alkaline Na 2 SO 4 type fluids associated with Fe-Mg phyllosilicates, zeolites, and gypsum. Alteration by S- and Cl-free fluids equilibrated with current atmospheric CO 2 leads to near-neutral low-salinity Ca-Na bicarbonate solutions at the top of the weathering profile (Figs. 1f and 4b). These compositions are consistent with chemistry of surface waters in cold basaltrich terrestrial regions (Gislason et al., 1996; Stefansson and Gislason, 2001; Pokrovsky et al., 2005). At depth, alkaline Na sulfate-chloride solutions dominate, though deeper gypsumsaturated solutions (W/R < 1, Fig. 4b) are Na-Ca chloride type. Magnesium is present in trace amounts (~10-8 mole per kg H 2 O) controlled by solubility of corresponding phyllosilicates. The data on concentrations of Fe-bearing solutes in low-water environments are restricted by the uncertain thermodynamics of Fe-phyllosilicates (Sections 4.1 and 4.3) Weathering by percolating solutions The progress of chemical weathering in an upper rock layer is seen in results of Model II where basalt reacted with parcels of H 2 SO 4 -HCl type solutions equilibrated with current atmospheric O 2 and CO 2. Figure 5 illustrates that a moderately altered assemblage with smectites, zeolites, and calcite is further transformed to the mineralogy where kaolinite, silica, 13

14 and goethite dominate. Earlier formed saponite is altered to montmorillonite. Alteration of montmorillonite by low-ph solutions leads to formation of kaolinite, which dissolves in new portions of acidic fluids. Gypsum forms at an intermediate stage of weathering and then dissolves. Only trace amounts of gypsum are present at advanced stages of alteration when the mineralogy is dominated by kaolinite and amorphous silica. Goethite is present among initially formed secondary minerals and it becomes the second major phase in strongly leached compositions. Minor alunite forms in kaolinite-bearing assemblages and rutile is present throughout. The bulk composition of altered rocks (Fig. 5b) indicates a limited leaching of Si, Fe, and Ti even in the contact with strongly acidic fluids. The limited leaching of Fe is accounted for by the formation of goethite in contact with O 2 -saturated solutions. The leaching ability increases in the sequence: Ti, Si, Fe 3+, Al, K, Mn, Ca, Mg, and Na. At early stages of alteration, the solution anions (SO 2-4 and Cl - ) are fully compensated by Na + leached from the rock and the ph of alteration solutions is slightly alkaline (Figs. 5c and 5d). As acidic weathering progresses, the early alkaline Na 2 SO 4 type solutions are replaced by neutral and acidic fluids dominated by Ca, Mg, Al, and then Fe 3+ sulfates. The abrupt changes in the solution composition and the ph reflect either appearance or dissolution of secondary phases. The transfer of cation-bearing fluids to lower rock layers decreases the neutralization capacity of forming acid-resistant minerals and the ph of alteration solutions approaches that of the initial fluid. The silica-goethite mineralogy of the strongly altered rock has a limited ability to affect solution ph in the modeled upper layer. Figure 6 shows typical weathering profiles calculated in the framework of Model II. The most altered surface layers are dominated by silica and goethite, and subsurface layers are rich in either montmorillonite or kaolinite. The oxidation of solution Fe 2+ to goethite by dissolved O 2 14

15 occurs only at near-surface conditions. The dissolved O 2 is fully consumed in upper rock layers and ferric phases do not form at depth where O 2 -free solutions allow the stable occurrence of ferrous sulfides (pyrite) and phyllosilicates. Deeper layers are rich in Fe-Mg clay minerals (saponite, chlorite, amesite) and contain zeolites, minor carbonates, and unaltered minerals of basalt. Gypsum forms in subsurface layers and is abundant in middle parts of profiles, especially within the Fe-Mg clay units. The upper Al-, Si-rich unit composes ~20 30% of the profile s thickness. The comparison of two profiles in Fig. 6 indicates changes in mineralogy as weathering progresses. Near-subsurface montmorillonite is replaced by kaolinite and amorphous silica, gypsum disappears in the Al-clay unit, upper saponite-rich layers are altered to Fe 2+ chlorite-bearing assemblages, zeolites are dissolved in the middle part of the profile, and pyrite is oxidized at deeper layers. Goethite, montmorillonite, gypsum, and zeolites form deeper, and the alteration front advances downward. The fluid percolation models demonstrate a significant leaching of Na, Mg, Ca, and Mn from upper rock layers (Fig. 7a). The uppermost layer mostly preserves Ti, Fe, and Si from basalt. The subsurface rock (layer 2) is enriched in Al and K relatively to basalt owing to formation of acid-resistant Al-rich clay minerals. Deeper alteration layers are strongly depleted in Na but mostly preserve other elements. The altered subsurface rocks and middle parts of the profile are enriched in S (up to 15 times) because of gypsum formation. These rocks also contain up to 50% more Mg than basalt owing to formation of phyllosilicates from Mg-bearing solutions percolated from above. The behavior of Mn is accounted for by dissolution of Mn-bearing silicates in acidic solutions and formation of carbonates at depth. The solution composition strongly varies with depth and reflects diverse solubilities of secondary phases. With depth, alteration solutions change from acidic sulfate compositions to 15

16 near-neutral MgSO 4 fluids. The formation of Mg sulfate compositions in middle parts of the profile (layers 3 6 in Fig. 7b) in association with Fe-Mg phyllosilicates is accounted for by the transfer of Mg 2+ from upper layers, a low solubility of Fe-, Na-, K-, and Ca-bearing secondary phases (phyllosilicates, zeolites, and gypsum) which do not provide sufficient cations to charge balance SO 2-4. As weathering progresses, MgSO 4 type solutions become dominant at deeper layers. In the ideal case of complete water-rock equilibration, Ca-Na sulfate-chloride alkaline compositions could form in lower parts of weathering profiles (see Section 4.3). The increase of ph with depth (Fig. 7c) reflects the consumption of H + through dissolution of minerals and an increasing neutralization capacity of lesser leached materials. The lower salinity of deep Ca/Na sulfate-chloride solutions (Fig. 7b) reflects the low solubility of coexisting gypsum, calcite, and Mg- and Ca-bearing silicates, and trapping of SO 2-4 in gypsum in upper layers Kinetic models The results obtained with Model III are roughly similar to outcomes of other models but constrain temporal fates of primary minerals and secondary phases formed at earlier stages of alteration. Figure 8 illustrates the weathering progress in a top layer of basalt affected by parcels of the H 2 SO 4 -HCl solution equilibrated with atmospheric CO 2 and O 2. Olivine dissolves faster than plagioclase and hypersthene, though relative dissolution rates vary with changing ph of alteration fluids. Plagioclase and diopside dissolve slowly and are present in strongly leached mineral assemblages. Early-formed secondary minerals (siderite, daphnite, gypsum) dissolve in new portions of acidic fluids. The advanced secondary mineralogy consists of amorphous silica, kaolinite, and goethite, which then dissolves. As alteration progresses, the volume of the altered 16

17 rock decreases owing to dissolution of primary and secondary phases which does not compensate the formation of hydrated phases with large molar volumes (Fig. 8b). A moderate ph of alteration solutions at early stages of weathering is accounted for by neutralization of acidic fluids through dissolution of primary minerals. Early solutions are nearneutral MgSO 4 type compositions (Figs. 8c and 8d) mainly accounted for by dissolution of olivine. At later stages, lower amounts of Mg, Ca, and Na in remaining minerals limited the supply of cations and the ph of formed alteration fluids decreases. The dissolution of plagioclase in moderately acidic solutions leads to a temporal Al-sulfate solution. Finally, a slow dissolution of the goethite-kaolinite-silica-diopside assemblage supplied Fe 3+ together with minor Ca, Mg, and Si in strongly acidic fluids which become similar to initial H 2 SO 4 -HCl-CO 2 compositions. Modeling of weathering profiles demonstrates that secondary precipitation in upper and middle units mainly occurs through dissolution of alteration products formed at earlier stages. Montmorillonite forms through dissolution of Mg-Fe phyllosilicates, and dissolution of other phyllosilicates precedes precipitation of kaolinite and amorphous silica. The most altered nearsurface layers are dominated by silica and kaolinite (Figs. 8 and 9). These phases accumulate because they dissolve slower than other secondary minerals in low-ph solutions. Other secondary minerals are serpentine, montmorillonite, Mg and Fe-Mg saponite, goethite, and gypsum. Amesite and Fe-chlorite are minor phases in the modeled three rock layers. Secondary minerals coexist metastably with unaltered plagioclase and pyroxene grains, though olivine is almost dissolved. The presence of serpentine could be accounted for by its slow dissolution after precipitation from rapidly altered olivine. The relatively low solubility and slow dissolution rates of gypsum account for its prompt precipitation from slightly acidic or neutral fluids and survival in evolving weathering profiles. As in this illustrative example, kinetic models demonstrate a 17

18 common presence of primary minerals and metastable secondary phases. They are present because they dissolve slower than the downward advance of the weathering front where fastdissolving phases (olivine) are most affected Freezing of alteration solutions and thawing of salt-ice mixtures A decrease of surface and subsurface temperature after the supposed formation of Noachian weathering profiles could have caused partial or complete freezing of alteration solutions. We modeled equilibrium and fractional freezing of representative fluids formed in the middle and lower parts of alteration profiles (Figs. 4a and 7b; Table 3) in the framework of Models I and II. Equilibrium freezing of near-neutral Mg-Na and Mg sulfate solutions formed in middle parts of weathering profiles leads to a major early precipitation of gypsum together with ice at -1 ± 0.5 C (Fig. 10). Gypsum is either first or second abundant salt down to eutectic at -32 C to -35 C. Depending on the initial solution composition, either mirabilite (Na 2 SO 4 10H 2 O) or meridianite (MgSO 4 11H 2 O) precipitate at and below ~ -4 C. Freezing down to ~ -15 C causes precipitation of majority of mirabilite and meridianite. Hydrated Na and Mg chlorides form at near-eutectic conditions. The sequential precipitation of Ca, Na, and Mg sulfates accounts for the compositional evolution of initial Mg-Na and Mg sulfate fluids toward Mg-Na chloride brines which compose only a small fraction of initial solution (< 0.03 at T < -15 C). At T = -4 ± 4 C, the relative concentration of Mg sulfate is higher than in initial solutions because of precipitation of Ca and, in some cases (Fig. 10a), Na sulfates. Therefore, a moderate freezing of sulfate solutions formed in the middle part of weathering profiles produces MgSO 4 -rich brines (Figs. 10b and 10d). These models also imply that warming of ice-salt assemblages up to ~ -5 ± 2 C 18

19 will lead to MgSO 4 -rich brines. Similar conclusions could be made based on calculations of fractional freezing of Mg-rich solutions (Fig. 11). Freezing of supposed Mg-poor solutions from lower parts of weathering profiles is accompanied by an early precipitation of gypsum and mirabilite (in some cases) and the formation of Na-Ca chloride brines (Fig. 12). An excess of Ca 2+ over SO 2-4 in the initial composition (Table 3, solution D) and/or precipitation of mirabilite from Na-rich fluids (solution C) causes accumulation of Ca 2+ which is not controlled by gypsum. A significant increase in Ca 2+ concentration in chloride brines decreases eutectic temperature. The precipitation of hydrohalite (NaCl 2H 2 O) below -22 C further contributes to the formation of Ca chloride brines which dominate from ~ -30 C to eutectic at ~ -52 C when antarcticite (CaCl 2 6H 2 O) precipitates together with hydrohalite and ice. The latter path agrees with the CaCl 2 -NaCl-H 2 O phase diagram and the freezing scenario discussed by Burt and Knauth (2003). The modeled freezing could give insights into formation of martian chloride brines and deposits (Osterloo et al., 2010). Figure 13 summaries major solution types formed through acidic weathering and subsequent freezing. 4. Discussion 4.1. Modeled mineralogy of weathering profiles The modeled weathering profiles demonstrate the formation of compositionally smooth vertical stratigraphies of dominated phyllosilicates, a thin Al-, Si-rich and oxidized upper layer, and relatively thick mineralogically diverse lower units with Mg-Fe phyllosilicates. Lastly formed secondary phases coexist with other alteration products and remnants of primary minerals. These features are similar to general characteristics of terrestrial weathering crusts on 19

20 basalts (Nesbitt and Young, 1989; Nesbitt and Wilson, 1992; Chamley, 1989; Velde and Meunier, 2008; Greenberger et al., 2012) and Noachian clay-bearing stratigraphies. Specific matches with martian profiles in the Mawrth Vallis region include the appearance of montmorillonite- and/or kaolinite-rich assemblages in upper units (Bishop et al., 2008), the presence of Fe 3+ oxide/oxyhydroxide phases in uppermost stratigraphic layers (McKeown et al., 2009; Farrand et al., 2009; Wray et al., 2008), the occurrence of a Fe 2+ phyllosilicate between the Al-rich clay and the smectite-rich units (Bishop et al., 2008), and a possible presence of zeolites (Michalski and Fergason, 2009; Bishop et al., 2013; Michalski et al., 2013). The similarity supports the assertion about formation of martian clay-bearing stratigraphies through alteration by percolating surface solutions. Our models demonstrate the formation of Al-Si phases through dissolution of earlier formed secondary minerals. This implies that the mineralogy of upper Al-Si-rich units may not reflect petrology of a protolith. At least some martian stratigraphies with Al-Si units could have formed through weathering of sedimentary deposits of Mg-Fe phyllosilicates. This notion is consistent with experimental (Altheide et al., 2010) and numerical models (Gainey et al., 2015; Zolotov and Mironenko, 2007b) for acidic alteration of phyllosilicates. In other places, the presence of pyroxenes and plagioclases in clay-bearing units (Poulet et al., 2008) and the rarely observed transitions between Fe-Mg phyllosilicate units and underlying country rocks (Carter et al., 2015) indicate mafic protoliths. The models give initial insights into weathering geochemistry rather than predict accurate mineralogy of alteration products. Used thermodynamic data for smectites (Wolery and Jove- Colon, 2004) are estimations, as well as recent data on Fe phyllosilicates (Catalano, 2013). Poorly crystallized phases with a variable composition are not considered, though they are 20

21 abundant in mafic martian sediments (Bish et al., 2013) and terrestrial weathering products under permafrost conditions (Pokrovsky et al., 2005). Models I and II do not consider metastable solids and all models do not include mixed-layers clay minerals such as smectite-chlorites, which are common on Earth and could be present on Mars (Milliken and Bish, 2010; Sun and Milliken, 2015). Models I and II assume complete water-rock equilibration which does not characterize low-temperature systems. As an example, the observed formation of Fe-chlorite should not be considered as a solid prediction. Although mixed-layer chlorite-smectites could form at early stages of basalt weathering (Pokrovsky et al., 2005), chlorites are rare minerals in weathering profiles of terrestrial mafic rocks (Chamley, 1989; Nesbit and Wilson, 1992; Velde and Meunier, 2008; Greenberger et al., 2012). Their occurrence is typically attributed to aqueous alteration of mafic rocks at elevated temperatures (> ~50 C; Nesbit and Young, 1989; Alt, 1999). The apparent presence of chlorites in exposed deep or uplifted crustal materials on Mars could be attributed to low-temperature hydrothermal processes (Ehlmann et al., 2011; Sun and Milliken, 2015). Modeling results with unsuppressed chlorites were revealed here because corresponding thermodynamic data (Holland and Powell, 1998) are based on multiple experiments and look more reliable than calculated data on Fe smectites. Runs with suppressed chlorites led to the formation of Fe-rich saponite, Fe-rich serpentine, stilbite, and magnetite in saponite-dominated lower parts of weathering profiles. It follows that Fe-saponite is more likely mineral than Fechlorite in weathering profiles of olivine-rich basalts. Our results for Fe-rich Adirondack martian basalt rather indicate the thermodynamic drive for formation a Fe-phyllosilicate, consistent with observations in the Mawrth Valley region (Bishop et al., 2008). The ambiguity related to the stability of Fe phyllosilicates affects the evaluation of solution composition (Section 4.3). 21

22 The presented results of Model III are fairly demonstrative and depend on uncertain parameters such as grain size and exposure of minerals to solution. Nevertheless, the general sequence of Mg-Fe-Ca-Na-K phyllosilicates, Al-rich smectites, kaolinite-group minerals, and amorphous silica seems valid Insights into Noachian weathering environments The models show that weathering by both H 2 SO 4 -HCl acidic and S-, Cl-free CO 2 -saturated fluids leads to roughly similar sequences of major clay minerals. The formation of similar mineralogical profiles requires different amounts of specific alteration fluids. The lower ph of initial fluids, the lower fluid mass is needed to develop weathering profiles. The formation of weathering profiles by slightly acidic solutions saturated with current atmospheric CO 2 requires a large amount of fluids (W/R > ~10 4 ), while strongly acidic H 2 SO 4 -HCl solutions produce comparable profiles of clay minerals at lesser amounts (W/R of ~10 2 ) (Fig. 1). The composition of initial solutions could be constrained from data on surface hydrology on ancient Mars. On the one hand, the occurrence of highly degraded impact craters and the valley networks in Noachian formations suggest surface water precipitation and a moderate fluvial erosion (e.g. Craddock and Howard, 2002; Mangold et al., 2012) which may not exclude weathering by large volumes of mildly acidic solutions. In the other hand, an alternative concept of mostly cold and dry Noachian epoch (e.g. Gaidos and Marion, 2003; Fairén, 2010) with brief impact-generated rainfalls (Segura et al., 2002), short-term eruption-related warming episodes (Tian et al., 2010; Wordsworth et al., 2013; Halevy and Head, 2014), and transient obliquity-related ice melting (Kite et al., 2013) could be consistent with weathering by smaller amounts of H 2 SO 4 -HCl type fluids. 22

23 The morphology of Noachian degraded craters and estimated erosion rates (~ mm/yr) imply arid climate with rare rainfall and surface runoff (Craddock et al., 1997). These erosion rates are similar to rates in terrestrial polar periglacial environments and are more consistent with lower masses of surface waters that are needed in high-ph weathering scenarios. The orientation of Noachian valley network (lack of correlation with topography) is consistent with arid climate (Luo and Stepinsky, 2012). Recent climate models which include geomorphology, variable obliquity, and occasional melting of surface ices (Wordsworth et al., 2013; Mishna et al., 2013; Kite et al., 2013) imply mostly sub-freezing Noachian surface environments. Greenhouse models for dense CO 2 -H 2 O atmospheres fail to achieve sufficient warming on early Mars, mainly because of CO 2 condensation in clouds and Rayleigh scattering of incoming sunlight (Kasting, 1991; Forget et al., 2013; Wordsworth et al., 2013). We advocate for the formation of late Noachian clay-bearing stratigraphies through widespread weathering of basaltic materials by moderate masses of initial H 2 SO 4 -bearing solutions generated by episodic volcanic degassing (Greeley and Craddock, 2009; Gaillard et al., 2013; Michalski and Niles, 2012) and impacts (Zolotov and Mironenko, 2007a). Acidic fluids could have affected surface materials through precipitation and/or partial melting of acid-bearing surface ices (Kite et al., 2013; Zolotov and Mironenko, 2007a). The acidic weathering pathway is also supported by the occurrence of Hesperian massive sulfate deposits, which could not have formed through water-rock interaction in that epoch (Section 4.5). Further constraints on Noachian weathering conditions could be obtained from non-clay mineralogy of supposed weathering profiles. Although major clay minerals in modeled profiles are similar, the composition and ph of initial solutions affect the occurrence of less abundant phases, such as carbonates, gypsum, and 23

24 amorphous silica. The modeled weathering by mildly acidic S-, Cl-free solutions equilibrated with low-co 2 atmospheres produces unabundant carbonates in middle parts of profiles and minute gypsum in lower parts, and could not produce silica phases (Figs. 1d, 2a, and 3). Weathering by solutions saturated with dense CO 2 atmospheres leads to formation of moderately abundant carbonates and amorphous silica in middle parts of the profiles and only trace gypsum in lower parts (Figs. 2b and 3). For high pco 2 cases (> ~0.15 bar), precipitation of abundant siderite is expected from partially neutralized solutions at depth. In contrast, alteration by H 2 SO 4 - HCl type solutions causes formation of abundant amorphous silica in upper parts of weathering profiles, moderate amounts of gypsum in the middle and lower parts, and minute carbonates in the lower parts (Figs. 1a, 5, 6, 8, and 9). The mineralogy of the Mawrth Vallis region (Section 1) is more consistent with weathering by H 2 SO 4 -bearing solutions. The acidic pathway is supported by the presence of amorphous silica at the top of the stratigraphic sequence in association with Al-rich clays (Bishop et al., 2008; 2013), by the presence Ca sulfates (Wray et al., 2010) and possibly other sulfates (Bishop et al., 2013; Michalski et al., 2013) in the middle and lower units, and by the lack of detection of carbonates. The unusually high thickness of supposed alteration sequence, even by terrestrial standards, supports an acidic scenario. Although amorphous Al-Si oxides such as allophane have not been considered in our models, they are likely coexisting with amorphous silica in deeply acid-leached materials. However, their probable presence in martian materials may not indicate low-ph weathering: allophane and other amorphous Al-Si compounds are common predicts of basalt weathering under terrestrial permafrost conditions (Pokrovsky et al., 2005). We observed the formation of Al- and K-bearing sulfate alunite in some acid-leaching models for the upper stratigraphic unit (Figs. 1a and 5a) but did not see jarosite. Other works 24

25 (e.g. Elwood Maden et al., 2004; Zolotov and Shock, 2005) show that jarosite forms at low W/R ratios from O 2 -bearing acidic sulfate solutions. The supposed high W/R ratios in upper parts of weathering profiles are not consistent with formation of jarosite. These results correspond to possible misidentification of martian jarosite with near-infrared and Mössbauer spectroscopy (McCollom et al., 2014). If the identification of jarosite is valid, its occurrence in upper stratigraphic units in the Mawrth Valley region (Farrand et al., 2009, 2014; Bishop et al., 2013) could be related to a subsequent acidic weathering of a newly-deposited basaltic material at low W/R ratios, consistent with the initial interpretation (Farrand et al., 2009). Some acid sulfate weathering in the Late Hesperian and Amazonian is consistent with the occurrence of amorphous silica and jarosite in the Valles Marineris region (Milliken et al., 2008) (Section 4.5). The formation of jarosite and other late-formed sulfates at low W/R ratios could be related to partial melting of acid-bearing surface ices and ice-dust deposits (Niles and Mishalski, 2009; Michalski and Niles, 2012; Kite et al., 2013; Zolotov and Mironenko, 2007a). Copiapite could have formed by the same way, though it was absent from our database. The scarcity of information about non-clay mineralogy of many other martian outcrops with stratified clays confine the evaluation of composition and masses of alteration fluids. The models (Figs. 1 3) do not exclude formation of some clays-bearing stratigraphies through alteration by S-, Cl-poor solutions with moderately acidic ph. However, the apparent widespread nature of Noachian weathering (Carter et al., 2015) suggests a common action of strong acidic solutions suggested for the Mawrth Vallis region. The models show that formation of Al-rich clays through leaching by S-, Cl-free solutions at low pco 2 (10-4 to 10-2 bar) conditions requites the W/R mass ratio of ~ (Figs. 1d, 2, and 3). These numbers agree with estimates for terrestrial deeply weathered basalts based on 25

26 chemical and isotopic data (Innocent et al., 1997; Benedetti et al., 2003). Although kaolinitegroup minerals form atop Mg-Fe smectite-rich layers in terrestrial profiles, sulfates, carbonates, and amorphous silica phases are either minor or not reported minerals (e.g., Greenberger et al., 2012). The W/R mass ratio of 10 4 implies that formation of a 1 m thick layer of kaolinite-group minerals requires leaching by a 30 km tall water column (assuming martian basalt density of 3 g cm -3 ). These numbers may not be consistent with the patchy appearance of Noachian valley networks, other observations, and climate models cited above. The apparent lack of abundant ferrous carbonates in martian clay-bearing stratigraphies do not support weathering in the presence of a dense (> ~0.5 bar CO 2 ) Noachian atmosphere, in accord with following evaluations. The tentative data on carbonate inventory in the crust are not consistent with a dense CO 2 atmosphere in the Noachian epoch (Niles et al., 2013; Edwards and Ehlmann, 2015). The Ar isotopic data on the ALH martian meteorite suggest an upper limit of 400 mbar CO 2 after ~4.2 Ga (Cassata et al., 2013). A dense CO 2 atmosphere could have not existed geologically long (~10 7 yr) due to the thermal escape (Tian et al., 2009; Lammer et al., 2013) and aqueous trapping to carbonates (Pollack et al., 1987; Manning et al., 2006; Kite et al., 2011). A replenishment of CO 2 in a dense atmosphere could have been limited because of the apparent incompletion of the atmosphere-crust CO 2 cycle without plate tectonics, though impact recycling of crustal carbonates could have played a role (Carr, 1999). A dense atmosphere could have not survived the impact erosion in the Noachian epoch (Melosh and Vickery, 1989; Zahnle, 1993; Manning et al., 2006). Lastly, a similar isotopic composition of carbon in carbonates from martian meteorites and the current atmosphere does not indicate much loss of atmospheric carbon since ~4 Ga (Webster et al., 2013). In other words, putative dense CO 2 atmosphere could have not existed during Noachian weathering. 26

27 4.3. The composition of alteration solutions The models show that formation of martian clay-bearing stratigraphies through weathering by H 2 SO 4 -HCl type solutions was accompanied by the formation of neutral and alkaline solutions enriched in Mg 2+, SO 2-4, Na +, Ca 2+, and Cl - ions and neutral aqueous complexes (MgSO 4, etc.). The dominance of these solutes in alteration fluids is also accounted for by the high solubility of Mg and Na salts: MgSO 4, MgCl 2, NaCl, Na 2 SO 4, and their hydrates. In contrast to Ca sulfates and carbonates, precipitation of these salts requites a significant increase in solution salinity through freezing or evaporation. Weathering by solutions with different initial ph causes preferential leaching of Na, K, Ca, and Mg from upper parts of weathering profiles (Fig. 1). Although the leaching of cations implies their transfer to lower layers, the composition of alteration fluids at depth differs from that of altering rocks owing to secondary precipitation and uneven dissolution of minerals (Figs. 7 9). Sodium and potassium are partially trapped in smectites, though a significant fraction of Na remains in solution (Figs. 4 and 7b). The concentration of Ca-bearing solutes is limited by precipitation of gypsum, carbonates, zeolites, and smectites. However, Ca 2+ could be among major cations in lower parts of weathering profiles (Figs. 4a and 7b). Leached Fe 2+ is not abundant in modeled fluids because of precipitation of low-solubility ferric oxides/oxyhydroxides in upper O 2 -bearing layers and ferrous phyllosilicates in anoxic layers. The calculated concentration of Fe 2+ species in anoxic layers is less reliable than data on other elements. The uncertainty is related to tentative thermodynamic data on Fe-bearing smectites and the ambiguity related to the formation of chlorite (Section 4.1). Models with the suppressed formation of chlorite show higher concentrations of Fe 2+ solutes (~1 to 100 μmole 27

28 per kg H 2 O) in equilibrium with Fe-saponite and Fe-Mg serpentine than in equilibrium with Fechlorite. Despite the formation of Mg-bearing smectites from local and transferred Mg species, Mg 2+ remains the major cation in middle parts of weathering profiles dominated by neutral MgSO 4 type solutions (Fig. 4a; Fig. 7b; Table 3). The elevated Mg 2+ concentration is accounted for by the necessity to compensate the abundant SO 2-4 ion supplied with initial solutions. As weathering progresses, neutral MgSO 4 type solutions occur at deeper depths in association with gypsum and Mg-bearing smectites (Figs. 6 and 7). In deeper environments, the modeled decline in Mg 2+ concentrations is accounted for by the consumption of SO 2-4 to gypsum and a decrease in SO 2-4 concentration at lower W/R ratios. Analogous Mg-deficient compositions characterize deep downward flows of seawater in suboceanic hydrothermal systems after trapping of oceanic SO 2-4 to anhydrite (German and Seyfried, 2014). In low-t martian conditions, the consumption of Mg from sulfate solutions could occur through ion exchange with Ca-bearing smectites followed by precipitation of gypsum (Vaniman et al., 2011; Wilson and Bish, 2011). However, the slow dissolution of Ca minerals in alkaline parts of martian profiles (Model III, Figs. 8 and 9) and stable occurrences of Ca-zeolites and diopside (Models I and II, Figs. 1 and 6) could limit the consumption of SO 2-4 to low-solubility Ca sulfates. The limited supply of Ca 2+ to solution could account for a metastable coexistence of MgSO 4 type fluids with partially altered basalts in deep parts of weathering profiles and below them. Alkaline Na-Ca sulfate-chloride solutions modeled for deep rock-dominated systems (Table 3) represent an ideal case of complete waterrock equilibration which may not be attained on Mars. The latter solutions could be less abundant than MgSO 4 type fluids owing to lower porosity of rocks at depth and the consumption of water through hydration in upper and middle parts of weathering profiles. 28

29 4.4. Fate of subsurface alteration solutions in the Noachian epoch The lack of detection of high-solubility sulfates and chlorides of Mg and Na in martian clay-bearing stratigraphies may indicate transfer of corresponding solutions away from weathering profiles. Even for supposed weathering by H 2 SO 4 -HCl solutions, the W/R mass ratio of ~10 2 (Fig. 1a) implies a removal of a large volume of neutralized fluids (~ m water column per 1 m of strongly leached rock with Al-clays). In the Noachian epoch, the migration of alteration solutions could have occurred through percolation deeper to the crust and a lateral subsurface movement. The migration could have been accompanied by salt precipitation through fluid freezing (Figs ), evaporation, and changes in solution composition through separation of brines from salts. Diverse pathways of subsurface migration and accumulation of neutralized solutions were possible depending on lateral and temporal variations in temperature, permeability, and reactivity of subsurface materials. The supposed widespread chemical weathering in equatorial and middle latitudes ( Ga, Carter et al., 2015; Ga in the Mawrsh Valley region, Loizeau et al., 2012) apparently occurred during an active phase of the Late Heavy Bombardment in the inner solar system (~ Ga, Gomes et al., 2005; Bottke et al., 2012). It implies a major effect of impacts on formation of acidic rains and aerosols (Segura et al., 2002; Zolotov and Mironenko, 2007a), finegrained sediments (Knauth et al., 2005), and the development of permeable megaregolith able to store alteration fluids. The low martian gravity makes the fragmented materials permeable down to km (Clifford et al., 2010). A high permeability of Noachian near-surface materials favored percolation of surface waters and limited lateral surface and near-subsurface migration of solutions and their accumulation in shallow surface depressions. The high permeability may 29

30 account for a limited appearance of sulfates in Noachian formations. In a case of a warm climate and/or a strong geothermal flow, alteration fluids could have filled pore spaces deep in the icefree equatorial megaregolith. In the likely scenario of a cold climate (Section 4.2), a significant faction of fluids could have accumulated at the upper boundary of low-permeability permafrost (Fig. 14). An uneven subsurface topography of that boundary could have caused a patchy distribution of fluids and coexisting salts. Subsequent fate of subsurface fluids could have been affected by cooling of the crust after an active phase of the Late Heavy Bombardment. Cooling led to partial or complete freezing of accumulated fluids which contributed to chemical (Figs ) and spatial separation of salts and brines. In one of possible pathways, a downward freezing of pore fluids further favored the accumulation MgSO 4 type brines at the upper boundary of the permafrost while some chloriderich solutions could have percolated to deeper layers. Some solutions could have released from overpressured aquifers formed in freezing environments (cf. Andrews-Hanna and Phillips, 2007). Precipitation of gypsum is expected from all low-salinity solutions at the beginning of freezing. Analogous precipitation could account for the formation of gypsum veins seen at the Meridiani Planum (Squyres et al., 2012) and Gale crater (Nachon et al., 2014). Mirabilite could be among early precipitates from Na-rich sulfate solutions (Figs. 10a and 12a). The deposition of Ca and Na sulfates through freezing of Mg-rich solutions enhanced the MgSO 4 -rich composition of fluids until after ~ -7 C when increasingly unabundant brines were presented by Mg-Na chlorides (Figs. 10 and 11). MgSO 4 -rich brines could be separated from precipitated Ca-Na sulfates during the downward freezing in weathering profiles and/or an upward advancing of the permafrost. Low-temperature (< -7 C) Mg-Na chloride brine could have percolated deeper into the permafrost living sulfates behind (Fig. 14). The separation could also be caused by migration 30

31 of MgSO 4 brines toward topographic lows of the advancing permafrost boundary. Precipitation of Ca and Na sulfates through partial freezing of presumed Mg-poor solutions led to formation of Na-Ca chloride brines (Figs. 12 and 13). Therefore, partial freezing could have contributed to further compositional deviation of probable Mg-rich and putative Mg-poor solutions formed in weathering profiles. Mg-Na-Ca sulfate compositions could be separated from Mg-Na, Na-Ca and Ca chlorides which are stable at lower temperatures. Evaporation models reveal roughly similar patterns of solution evolution. Although the supposed divergence of Mg sulfate and chloride-rich solutions may account for the observed appearance of corresponding salts (Gaillard et al., 2013; Osterloo et al., 2010), formation pathways of exposed salt-rich deposits require further considerations. Subsequent fate of subsurface alteration solutions and salts could have been affected by tectonic deformations, magmatism, rifting, ground ice melting, and changing groundwater levels that characterized the Hesperian epoch An origin of Hesperian layered sulfate deposits The Hesperian geological formations are characterized by massive deposits of layered sulfates seen in the Valles Marineris trough system and related depressions, in Arabia Terra, within Aram and other chaos terrains, and at the sources of the outflow channels in chaos regions (Gendrin et al., 2005; Glotch and Christensen, 2005; Glotch and Rogers, 2007; Murchie et al., 2009a, 2009b; Carter et al., 2013). Although layered sulfate deposits rarely appear in a close proximity to clay-bearing phyllosilicates, they co-occur within many low-latitude regions (DeLeit et al., 2012; Michalski et al., 2013). The thickness of layered sulfate-bearing deposits in chasmata reaches ~5 km (Lucchitta et al., 1994; Fueten et al., 2014). In chasmata and other deep 31

32 depressions, layered sulfates are located stratigraphically below phyllosilicates (DeLeit et al., 2012). In the Valles Marineris trough systems, layered sulfate deposits occurs at different elevations and often reveal deformation, folding, faulting, dipping, and erosion on deposits possibly formed in ancestral basins (Lucchitta et al., 1994; Schultz, 1998; Flueten et al., 2008, 2010, 2014; Bishop et al., 2009). Massive layered sulfate deposits are dominated by Mg sulfates, though Fe 2+ and Ca sulfates, and Fe 3+ oxides (hematite, nanophases) are reported as well (Gendrin et al., 2005; Le Deit et al., 2008; Mangold et al., 2008; Bishop et al., 2009; Murchie et al., 2009b; Roach et al., 2010a, 2010b; Flahaut et al., 2010; Fueten et al., 2010, 2014; Weitz et al., 2008, 2015; Christensen et al., 2001). In many cases (Mangold et al., 2008; Murchie et al., 2009b; Bishop et al., 2009), silicates are not reported as a major component of layered lighttoned sulfate deposits, though spectral signatures of Al- and Si-bearing minerals could be masked or shifted in sulfate-rich rocks (Noe Dobrea et al., 2012; Fueten et al., 2014). The sulfate-rich mineralogy indicates a water-related sedimentary origin. Although diverse formation scenarios have been discussed for massive sulfate deposits, probable pathways suggest precipitation in lakes (Nedell et al., 1987; Komatsu et al., 1993; Malin and Edgett, 2001; Quantin et al., 2005; Roach et al., 2010b; Lucchitta, 2010; Fueten et al., 2014), spring, playa, or cementation of pre-existing eolian deposits through evaporation of released salty groundwater (Andrews-Hanna et al., 2007, 2010; Andrews-Hanna and Lewis, 2011; Rossi et al., 2008; Mangond et al., 2008; Murchie et al., 2009b; Roach et al., 2010a). Deposition of sulfates could have been driven by repeated evaporation (Murchie et al., 2009b) and/or freezing of surface waters. One of alternative scenarios suggests the formation of sulfate deposits in chasmata and some other depressions through low-w/r acidic alteration within an airborne mixture of ice, mineral dust, and sulfuric acid during periods of high obliquity (Michalski and Niles, 2012). This 32

33 mechanism is suitable for formation of thin, draping, sulfate-silicate, jarosite-bearing deposits along wallrocks of canyons (Weitz et al., 2015) and could explain the formation of sulfatesilicate deposits in the Meridiani Planum (Niles and Michalski, 2012). In other places (e.g. in Gale crater, Le Deit et al., 2013), sulfate-bearing deposits may not be related to a release of groundwater water either. The presence of sulfate deposits that drape preexisting topography, rather than concentrating on topographic lows, places constraints on groundwater models. What was the chemical origin of sulfate sulfur in massive deposits? A supply of acid sulfates by Hesperian volcanic emanations (Bibring et al., 2006) may not be consistent with the lack of comparable volumes of Al-, Si-rich minerals, which should have formed through acidic leaching of cations and formation of abundant salts. Likewise, a widespread mobilization and accumulation of sulfates by Hesperian surface waters is inconsistent with geological data. Geological observations and hydrological evaluations indicate a drastic (several orders of magnitude) degrease in rates of surface erosion at the Noachian-Hesperian transition (Golombek and Bridges, 2000; Craddock et al., 1997; Fasset and Head, 2008). In contrast to Noachian degraded craters and valley networks (Craddock and Howard, 2002), Hesperian surfaces indicate a limited erosion by surface waters (e.g. Carr and Head, 2010). The abrupt change in the crater degradation style at the Noachian-Hesperian boundary implies corresponding changes in climatic conditions (Mangold et al., 2012). Valley networks in the cratered highlands become sharply unabundant in the Early Hesperian regions where signs of phyllosilicate-type weathering disappear as well (Fasset and Head, 2008). Hesperian valley drainage networks suggest very localized and episodic activity of liquid water typically related to recharge of subsurface solutions (Fasset and Head, 2008). It could not be enough surface aqueous solutions in the Hesperian epoch to account for weathering-related leaching of tremendous amounts of Mg 33

34 sulfate seen in massive deposits (Michalski and Niles, 2012). However, Hesperian subsurface hydrological processes should have affected salts and brines formed in the Noachian epoch. Geological observations and orbital mineralogical data indicate a localized subsurface mobilization and release of compositionally-diverse crustal fluids in the Hesperian epoch. These processes were related to a widespread volcanism and intrusive magmatism, tectonic events linked to the formation of Valleys Marineris depressions, regional changes in the temperature of the crust, melting of ground ice, and changing in the groundwater level (Carr, 1996; Carr and Head, 2010; Andrews-Hanna and Lewis, 2011). The outflow channels, pitted and chaos terrains are indicators of massive ground ice melting (e.g. Carr, 1996; Clifford and Parker, 2001). The formation of some valleys and deltas is accounted for by a short-lived local mobilization of water from the cryosphere (Hauber et al., 2013). In some cases, the release of solutions could be related to bursts of pressurized ice-sealed aquifers (Carr, 1979; Andrews-Hanna and Phillips, 2007). In addition to common releases of low-salinity waters through rapid ice melting, a supply of salty groundwater was suggested for several Hesperian craters (Wray et al., 2011) and for the Valley Marineris canyon system (Murchie et al., 2009b; Andrews-Hanna and Lewis, 2011). The presence of Fe 3+ oxides/oxyhydroxides in sulfate deposits is commonly attributed to aqueous oxidation of Fe 2+ by dissolved atmospheric O 2. Similar processes could have occurred at Aram Chaos and other chaos terrains in which layered sulfate deposits with hematite were detected (Glotch and Rogers, 2007; Liu et al., 2012). In other places (e.g. Meridiani Planum), formation of sulfate-bearing deposits may (Arvidson et al., 2006; Andrews-Hanna et al., 2007, 2010) or may not (Niles and Michalski, 2009) be related to release of subsurface waters. Here we suggest a formation of some massive layered sulfate deposits through a Hesperian remobilization of subsurface sulfate-rich solutions and salts formed during the widespread 34

35 chemical weathering in the Noachian epoch (Fig. 14). The release of Noachian salts in the Hesperian epoch explains the apparent deficiency of salts in Noachian deposits, which was first noted by Milliken et al. (2009). The remobilization could have involved partial melting of icesalt mixtures within warming permafrost which became more permeable for solutions. In some cases, subsurface fluid migration could have been related to tectonically-driven changes in topography of the permafrost. Subsurface solutions slowly migrated to newly-formed topographic lows such as chasmata in the Valles Marineris system, deep fresh craters (Wray et al., 2011), and depressions caused by through melting of ground ice. Salt deposits accumulated in depressions through water freezing and/or evaporation together with some airborne silicate dust. In the Valles Marineris trough system, initial deposition took place in ancestral basins and often occurred synchronously with the subsidence of early depressions (Lucchitta et al., 1994; Fueten, et al., 2008, 2010, 2014). These processes were followed by a major opening of troughs, and structural deformation, erosion, and redeposition of sulfate-rich deposits. This scenario explains the occurrence of dipped deposits on different topographic levels. Some late sulfate deposits formed through freezing and/or evaporation of groundwater released on chasma floors (e.g. Murchie et al., 2009b; Roach et al., 2010b). The supposed widespread Noachian weathering (Le Deit et al., 2012; Carter et al., 2015) corresponds to the large volume of Hesperian sulfate-bearing deposits in the Valles Marineris system. The volume of sulfate deposits in chasmata (Michalski and Niles, 2012) implies remobilization of subsurface fluids and salts from a large surface area. The common exposure of Noachian clay-beading stratigraphies on plateaus, chasma walls, pit walls, and impact crater rims in the Valles Marineris region (Le Deit et al., 2012) supports this notion. The suggested subsurface remobilization (Fig. 14) and 35

36 deposition of salts in shallow ancestral basins may not require a rise of groundwater layer in icefree subsurface conditions modeled by Andrews-Hanna et al. (2010). A rapid local ground ice melting followed by formation of outflow channels could have not produced much sulfate deposits from low-salinity water outflows. This notion agrees with the apparent scarcity of salt deposits associated with outflow channels and especially their sink areas, such as Northern Plains. In Aram, Aureum, and Iani Chaos, formation of sulfate-bearing deposits could be related to a slow release of salty groundwater to depressions formed through a rapid ground ice melting and the release of subsurface water (Glotch and Rogers, 2007; Andrews-Hanna and Lewis, 2011). The appearance of sulfates at sources of outflow channels could be explained by filling of topographic lows by salty groundwater after major releases of low-salinity waters. The apparent dominance of MgSO 4 hydrates in Hesperian sulfate deposits is consistent with a subsurface lateral migration of corresponding sub-freezing (down to a few degrees below 0 C) brines partially separated from less soluble Na- and Ca-sulfate compositions. However, only a partial separation of MgSO 4 -rich solutions in modeled thawing/freezing processes (Figs. 10 and 11) implies a presence of sulfates and chlorides of Ca and Na in forming deposits. The presence of Ca sulfates in lower stratigraphic units in some layered sulfate deposits (Weitz et al., 2015) could be accounted for by their deposition at early stages of evaporation or freezing of groundwater-fed lakes (Kuzmin et al., 2009). The expected presence of Na sulfates and chlorides in massive sulfate deposits remains to be confirmed. Our basic scenario suggests deposition of sulfate-rich salts from released neutral MgSO 4 type waters. This pathway does not exclude subsequent actions of acidic solutions. The occurrence of amorphous hydrated silica and jarosite within the Valles Marineris chasmata and 36

37 on surrounding plateaus could be related to post-depositional acid sulfate weathering in the late Hesperian to Amazonian in water-deficient environments (Milliken et al., 2008; Roach et al., 2010b; Weitz et al., 2015). The supposed late formation of jarosite in the Mawrth Valley (Farrand et al., 2009) and other regions agrees with this interpretation. The separate occurrence of jarosite-bearing deposits in Ophir Chasma may reflect sub-glacial actions of acidic solutions after formation of MgSO 4 -rich deposits (Cull et al., 2014). In Ius Chasma, the presence of jarosite and amorphous silica in association with some Mg sulfate deposits (Wendt et al., 2011) could be related to acidic alteration by the same late process that caused formation of jarosite and silica on surrounding plateaus and within other chasmata (Milliken et al., 2008; Bishop et al., 2009; Metz et al., 2009; Roach et al., 2010b). In some chasmata, jarosite-bearing deposits on elevated topographic levels and wallrocks could have formed through airborne precipitation of acid-bearing ices or aerosols followed by in situ alteration (Wendt et al., 2011; Michalski and Niles, 2012; Weitz et al., 2015), percolation of low-t acidic brine pockets in ice toward underling rocks (Zolotov and Mironenko, 2007a), and a transient melting of ice (Kite et al., 2013). However, the proposed airborne formation of all layered sulfate deposits in chasmata (Michalski and Niles, 2012) needs reconciliation with signs of aqueous deposition of salts at lower topographic layers in Ius, Candor, and Melas Chasmata (Murchie et al., 2009b; Roach et al., 2010b; Weitz et al., 2015) and remains to be supported by a definitive detection of Al-Si minerals (cf. Roach et al., 2010b; Fueten et al., 2014) and acidic sulfates (jarosite) in massive MgSO4-rich formations. Note that a coexistence of sulfates with Al-Si minerals could be explained by infiltration of preexisting eolian silicate deposits by released groundwater (e.g. Murchie et al., 2009b) and by aqueous co-deposition of sulfates and the airborne silicate dust in depressions. 37

38 5. Concluding remarks The decent similarity of modeled basaltic weathering profiles with martian clay-bearing stratigraphies agrees with a widespread chemical weathering of mafic Noachian rocks. Weathering by low-ph H 2 SO 4 -bearing solutions is more consistent with observations than alteration by S-free fluids. Weathering by acidic sulfate solutions does not require elevated water/rock ratios, a warm climate or a dense CO 2 atmosphere. It leads to formation of Ca sulfates in middle parts of weathering profiles together with neutralized Mg-rich sulfate solutions at depth. Weathering by slightly acidic fluids equilibrated with current atmospheric CO 2 requires large volumes of water which could be inconsistent with geological observations and climate models. Weathering by solutions equilibrated with a dense CO 2 atmosphere leads to formation of abundant subsurface carbonates, which are not observed in clay-bearing stratigraphies. A plausible formation scenario for Noachian clay-bearing stratigraphies and some Hesperian sulfate deposits is as follows. A vast majority of martian phyllosilicates and salts could have formed in Noachian by initially low-ph solutions at low pco 2 atmospheric conditions (< ~0.15 bar) followed by subsurface aqueous remobilization of high-solubility salts in subsequent epochs. Noachian weathering was caused by transient volcanic- and impact-generated H 2 SO 4 -HCl acidic rainfalls during the Late Heavy Bombardment and affected much of equatorial and middle latitudes. Weathering was facilitated by transient impact-, volcanism-, and obliquity-related warming and surface ice melting, consistent with climate models (Wordsworth et al., 2013; Mischna et al., 2013; Kite et al., 2013; Halevy and Head, 2014). Leaching of cations from upper layers of permeable mafic materials led to formation of hydrated Al-Si clay, Fe and Ti oxide, and 38

39 amorphous phases. At depth, rock alteration by percolated solutions led to precipitation of Fe- Mg phyllosilicates, Ca sulfates, zeolites, and minor carbonates. Some minerals of basalts (pyroxene, plagioclase) survived in middle and lower parts weathering profiles. The formation of stratified clay-bearing assemblages through weathering by initially acidic fluids produced neutral and alkaline Mg-, Na-, Ca-bearing sulfate-chloride solutions. Nearneutral MgSO 4 type solutions formed in middle parts of weathering profiles. The models suggest formation of alkaline Na-Ca sulfate-chloride solutions in lower parts. The formation of latter solutions could have been limited by incomplete dissolution of Ca minerals, which restricted trapping of SO 2-4 ion to low-solubility Ca sulfates. In such a case, metastable MgSO 4 type solutions occurred in lower parts of weathering profiles and beneath them. Subsequent chemical evolution of alteration solutions was affected by temperature, humidity, and permeability of the subsurface and included freezing, evaporation, migration, and melting of ice-salt assemblages. Freezing of MgSO 4 type solutions led to precipitation of Ca, Na, and Mg sulfates and formation of Mg-Na chloride brines stable down to down to ~ -35 C. Freezing of putative Na-Ca sulfate-chloride solutions caused formation of Na-Ca chloride and then Ca chloride brines existing down to -52 C. The precipitation of Ca-Na-Mg sulfates, a deep percolation of chloride brines, and a lateral migration of solutions contributed to compositional and spatial diversity of brines and salts in the subsurface. The diversity could have affected the composition of salts precipitated from solutions released in the Noachian and subsequent epochs. In the Hesperian epoch, major tectonic deformations and magmatism caused local melting of the permafrost and remobilization of subsurface salts and solutions formed in the Noachian epoch. Apparently abundant near-neutral MgSO 4 type fluids drained toward surface depressions. Sulfate deposits formed through freezing and/or evaporation of released subsurface solutions in 39

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56 Table 1 The composition of olivine basalt. Bulk composition mole kg -1 Normative mineralogy wt % O Hyperstene 18.9 K Diopside 13.2 Mg Plagioclase, An Ca Olivine, Fo Al Magnetite 5.3 Si Ti S Na Mn Fe Note: The data correspond to a simplified composition of the Adirondack-type olivine martian basalt (McSween et al., 2006). The used normative mineralogy has no S-, Cl-, and Mnbearing phases. 56

57 Table 2 Numerical approaches to model weathering profiles. Model I II III Description Equilibrium calculations in closed water-rock systems at variable W/R ratio Equilibrium calculations in closed water-rock-gas systems in a series of rock layers interacted with a sequence or water solution parcels percolated from above The same as for Model II but with a consideration of rates of mineral dissolution 57

58 Table 3 Concentration of major solutes in typical solution types modeled for middle (A and B) and lower (C and D) parts of weathering profiles, mole (kg H 2 O) -1. Species MgSO 4 type Na/Ca sulfate-chloride type A B C D SO Mg o MgSO Cl Ca o CaSO Na ph Eutectic, C Note: Solutions (A) and (C) are referred to W/R = 40 and 20 (Fig. 4a), respectively. Solutions (B) and (D) correspond to layers 6 and 7 (Fig. 7b), respectively. The solutions form through weathering of basalt by H 2 SO 4 -HCl solutions at 0 C and 1 bar. 58

59 Figure captions Fig. 1. The equilibrium secondary mineralogy, leaching pattern of elements, and solution ph in water-basalt type closed systems at variable W/R ratio at 0 C and 1 bar. Plots (a) (c) are for the initial H 2 SO 4 -HCl acidic solution and plots (d) (f) are for the initial S-, Cl-free fluid. Both solutions are equilibrated with current atmospheric 5.3 mbar CO 2 and 7.6 μbar O 2. The calculations are performed with Model I. The decreasing W/R ratio could be a proxy of the depth in a weathering profile. Alternatively, an increase in W/R ratio gives insights into progress of chemical weathering at a certain depth. In (a), minor phases are alunite (W/R > ~135), rutile (W/R > ~25); titanite, lawsonite, analcime, calcite, rhodochrosite (W/R < 25), and pyrite. Goethite forms at higher W/R ratios than shown. In (d), minor phases are rutile, titanite, pyrite, rhodochrosite, calcite, and gypsum (W/R < 1); gibbsite replaces kaolinite at W/R > Water/Rock mass ratio Water/Rock mass ratio diopside gypsum stilbite calcite lawsonite kaolinite stellerite Mineral volume, cm 3 per 1 kg of initial rock stellerite analcime kaolinite diopside stilbite amorph. SiO 2 montmorillonite Fe-chlorite saponite montmorillonite saponite Fe-chlorite Mineral volume, cm 3 per kg of initial rock a b c d Y Data Y Data Na Mg K Ca Fe Al Si Ti 0.1 Composition relatively to inital rock 1 Fe K K Mg Ca Al Si Ti Na Mn Fe Mg Mn Na Ca Composition relatively to inital rock Si Al, Ti Al Fe Mn K Ti e W/R mass ratio W/R ratio Depth ph ph Depth Alteration progress f Alteration progress 59

60 Fig. 2. The equilibrium secondary mineralogy and solution ph in water-basalt type closed systems at the variable W/R ratio at 0 C. The calculations are performed with Model I to explore weathering by putative CO 2 -rich solutions on ancient Mars. Initial S-, Cl-free aqueous solutions were equilibrated with atmospheric 7.6 μbar O 2, and with 0.1 bar (a) and 1 bar CO 2 (b and c). In (a), minor phases are titanite (W/R < 200), rutile (W/R > 200), rhodochrosite (W/R = ), gypsum (W/R < 0.6), and pyrite. In (b), minor phases are titanite (W/R < 18), rutile (W/R > 18), rhodochrosite (W/R = ), gypsum (W/R < 0.6), and pyrite. Water/Rock mass ratio am. SiO 2 calcite analcime lawsonite stellerite stilbite kaolinite montmorillonite diopside 0.1 bar CO 2 saponite Fe-chlorite W/R mass ratio calcite analcime lawsonite kaolinite amorph. SiO 2 stilbite diopside Mineral volume, cm 3 Mineral volume, cm per kg of initial rock per kg of initial rock 3 stellerite siderite 1 bar CO 2 montmorillonite saponite Fe-chlorite W/R mass ratio bar a b c Depth CO ph Alteration progress 60

61 Fig. 3. The selected mineralogy of modeled weathering profiles formed at dense CO 2 atmospheres. The data are for closed water-basalt system (Model I) in which initial S-, Cl-free fluids are equilibrated with atmospheric O 2 (~10-5 bar) and variable pco 2 to represent assumed ancient martian environments at 0 C. The black curves correspond to equal volumes of certain phyllosilicates (see Fig. 2 for the complete mineralogy for 0.1 bar and 1 bar CO 2 ). The grey and dotted curves mark the fields where amorphous silica and siderite form, respectively kaolinite Water/Rock mass ratio montmorillonite saponite amorphous silica siderite Depth 10 1 Mars now CO 2 partial pressure, bars 2 61

62 Fig. 4. The speciation of aqueous fluids formed through aqueous alteration of basalt at 0 C. The data correspond to complete chemical equilibration in closed water-rock type systems in the framework of Model I. In (a), basalt interacts with acidic H 2 SO 4 -HCl solution; other data are shown in Figs. 1a, 1b, and 1c. In (b), basalt reacts with S-, Cl-free solution; other results are shown in Figs. 1d, 1e, and 1f. The concentration of Cl - corresponds 0.1 wt% Cl in martian basalt. The decrease in concentrations of Ca- and sulfate-bearing species at W/R < 1 is related to precipitation of gypsum. Water/Rock mass ratio Ca 2+ Cl - Fe 2+ a 10 5 b Ca 2+ CaSO 4 2- SO 4 Mg 2+ MgSO 4 Na + - NaSO HCO NaHSiO Mole per kg H 2 O Mole per kg H 2 O W/R ratio SO Cl Ca CO CaSO NaSO Na

63 Fig. 5. The major minerals (a), bulk composition of solids (b), solution composition (c), and fluid s ph (d) which represent weathering of an upper basalt layer by H 2 SO 4 -HCl solutions. The calculations were performed with Model II in which 1 kg of basalt sequentially reacted Volume, cm 3 per kg of initial rock calcite Surface rock layer W/R = 1, open system montmorillonite stellerite saponite kaolinite gypsum amorphous silica goethite Solution parcels a with solution parcels (1 kg per parcel). The system was open with respect to current martian atmospheric CO 2 and O 2 at 0 C and 1 bar. In (a), minor phases are pyrolusite (parcels < 140), alunite (parcels ), and rutile. In (c), sulfur corresponds to the sum of SO 2-4, MgSO4 0, CaSO 0 4, and NaSO - 4 concentrations. Carbon is in HCO - 3, CO 2-3, and dissolved CO 2. In strongly acidic fluids (parcels > 130), Fe is in ferric species. ph Mole per kg H 2 O Mole per kg of initial rock O H Si Fe Fe, H Ca Mg S Al C Ti Na Mn K Solution parcels Solution parcels Na Alteration progress S (in sulfates) Mg Al Ca Cl (Cl - ) Fe Mn Si Na C K K Solution parcels b c d 63

64 Fig. 6. The modeled equilibrium mineralogy of weathering profiles after percolation of 175 (a) and 210 (b) parcels of H 2 SO 4 -HCl fluids through dimensionless basaltic layers. The data correspond to water-rock equilibration within each layer at 0 C and 1 bar in the framework of Model II when the rock layers (1 kg each) was affected by 250 parcels of fluids (1 kg each). The initial fluids were equilibrated with atmospheric 7.6 μbar O 2 and 5 mbar CO 2. In (a), minor minerals are stilbite (layers > 3), rutile (layers 1 6), titanite (layers > 7), pyrite (layers > 3), rhodochrosite (layers > 3), and calcite (layers > 7). In (b), minor minerals are rutile (layers 1 7), pyrite (layers > 4), and rhodochrosite (layers > 3). Lock layers Rock layers After 175 solution parcels After 210 solution parcels goethite amorphous SiO 2 kaolinite montmorillonite Fe-chlorite am. SiO 2 gypsum saponite amesite diopside stilbite a cm3/kg Volume of minerals, cm 3 per kg of initial rock goethite gypsum montmorillonite amesite amorph. SiO 2 kaolinite Fe-chlorite saponite Depth b 64

65 Fig. 7. The composition of altered rocks relative to initial basalt (a), speciation of alteration solutions (b), and the solution ph at a certain stage of weathering by percolating acidic solutions. The data correspond to the percolation of 175th solution parcel through rock layers (Fig. 6a). The ionic strength of solutions deceases with depth (from 0.15 at layer 1 to 0.06 at layer 8). In (a), the elevated Mg and Ca contents in altered rocks reflect the formation of phyllosilicates from elements leached from upper layers. The enrichment in S is accounted for by the formation of gypsum. In (a) and (b), the decrease in concentrations of rock s S, Mg 2+, and SO 2-4 below level 6 reflects the consumption of S to gypsum in upper layers. Analogous data for the solution parcel 210 demonstrate elevated concentrations of rock s S (in gypsum), Mg 2+ and SO 2-4 in solution, and a neutral ph down to level 8. The results for parcels 175 and 210 indicate percolation of MgSO 4 type solutions to depth as weathering progresses Mg Fe Ca Si, Ti Al, K O a CO 2 SiO 2 Cl - Fe 2+ MgSO 4 b c Rock layers 4 5 Na Mn S 4 5 MnSO 4 Mn 2+ HCO 3 - Mg Composition relatively to initial rock 6 7 NaSO 4 - Y Data 8 2- SO 4 Na + Ca 2+ CaSO Mole per kg H 2 O Y Data ph 65

66 Fig. 8. Chemical weathering of an upper layer of basaltic sand altered by parcels of H 2 SO 4 - HCl type solution equilibrated with the martian atmospheric CO 2 and O 2. In (a) and (b), secondary mineralogy is shown together with dissolving primary minerals. In (c) and (d), the composition and ph of alteration fluids are shown. In (c), Fe 2+ solutes dominate in alkaline and moderately acidic solutions and Fe 3+ dominates in acidic fluids (parcels >12). The data correspond to results obtained with the Model III in which 1 kg of rock reacts with 20 parcels of initial solutions, 10 kg each, the grain size of initial minerals is 0.1 mm in diameter, the grain size of secondary minerals is 1 μm, the degree of mineral exposure to solution is unity, and the allowed reaction time with each parcel of solution is 100 years. Volume, cm 3 per kg of initial rock Volume, cm 3 Mole per kg H 2 O ph 100 hyperstene amorphous SiO 2 plagioclase, An siderite C Ca Fe olivine daphnite Mg Solution parcels Na Si S (sulfate species) Al goethite gypsum montmorillonite magnetite Solution parcels Solution parcels kaolinite diopside Primary minerals Initial acidic solution Total volume X Data Secondary minerals C Cl (Cl - ) Fe a b c d 66

67 Fig. 9. The secondary mineralogy (a) and remaining primary minerals (b) in the basalt weathering profile calculated with the kinetic-thermodynamic approach (Model III). The profile has formed after percolation of 20 parcels of initial H 2 SO 4 -HCl type solution saturated with respect to atmospheric CO 2 and O 2. Each rock layer was allowed to react with each parcel of solution during 100 years and the total weathering time is 2000 years. Other parameters of the model are described in Fig. 8. In (a), minor gibbsite (< 2 cm 3 ) is present in layer 1; minor layer-3 minerals are Mg-Fe saponite, amesite, and Fe-chlorite. In (b), magnetite and minor olivine are present in layer 3 and minor diopside is present in layers montmorillonite Secondary phases a Rock layers goethite Fe-Mg serpentine kaolinite Mg-saponite gypsum amorphous silica vol Primary phases plagioclase (An 52, bytownite) hyperstene Depth Depth Mineral volume, cm 3 per kg of initial rock b 67

68 Fig. 10. The composition of solids and brines formed through equilibrium freezing of typical Mg-rich sulfate solutions from middle parts of weathering profiles. In (a) and (b), the data correspond to compositional changes of solution A in Table 3; in (c) and (d) the data are referred to solution B in Table 3. The dashed lines correspond to near-eutectic conditions. The plots could also be interpreted in terms of thawing of ice-salt mixtures. -1 ice Equilibrium freezing of solution A NaCl*2H O, hydrohalite 2 CaSO *2H O, gypsum 4 2 a 10-1 ice Equilibrium freezing of solution B MgSO *11H O 4 2 c log 10 Mole per kg H 2 O Moles per kg H 2 O MgSO 4 -rich MgSO *11H O, meridianite 4 2 Na SO *10H O, mirabilite Temperature, o C Cl - Mg 2+ Na + b Moles of salts per kg H 2 O MgSO 4 -rich CaSO *2H O 4 2 MgCl 2 *12H 2 O Na SO *10H O NaCl*2H O 2 Temperature, o C Cl - Mg 2+ Na + d Ca SO Temperature, o C Ca 2+ Moles per kg H 2 O -10 SO Temperature, o C

69 Fig. 11. The composition of solids and brines formed through fractional freezing of typical solutions from middle parts of weathering profiles. In (a) and (b), the data correspond to initial solution A in Table 3; in (c) and (d) the data are referred to solution B. In (a) and (c), the data show molar amounts of solids formed during incremental freezing with 0.5 C temperature step. Eutectic temperatures are reached below -30 C (Fig. 10). Fractional freezing of solut. A -1 ice gypsum -2 mirabilite a Fractional freezing of solut. B ice a gypsum meridianite c log 10 Mole per kg H 2 O Moles per kg H 2 O meridianite Temperature, o C MgSO 4 -rich Mg 2+ Cl - Na + Ca SO b Moles of salts per kg H 2 O Temperature, o C Temperature, o C MgSO 4 -rich Na + Ca Moles per kg H 2 O mirabilite Mg 2+ Cl - SO 4 2- Temperature, o C -10 d 69

70 Fig. 12. The compositional changes during fractional freezing of solutions from lower parts of weathering profiles. In (a) and (b), data correspond to initial solution C in Table 3. In (c) and (d), the data are for solution D. The dashed lines correspond to eutectic. Equilibrium freezing of these solutions demonstrates similar trends with an early precipitation of Ca and Na sulfates and the formation of Ca chloride brines below -30 C. Moles per kg H 2 O Moles of salts per kg H 2 O CaSO 4 *2H 2 O Na 2 SO 4 *10H 2 O Na + Cl- Ca 2+ SO 4 2- Fractional freezing of solution C ice Temperature, o C NaCl*2H 2 O, hydrohalite -30 CaCl 2 *6H 2 O, antarcticite -40 Na a b Moles of salts per kg H 2 O CaSO 4 *2H 2 O Moles per kg H 2 O Fractional freezing of solution D ice Temperature, o C CaCl 2 *6H 2 O NaCl*2H O Cl - Ca 2+ Na c d Temperature, o C Temperature, o C

71 Fig. 13. Major types of subsurface aqueous solutions formed through acidic weathering of martin basalt and partial freezing of alteration fluids. The left-hand side boxes show changes in solution composition in weathering profiles. Chloride-rich solution types are shown by ovals. The arrows with numbers are referred to precipitation of salts which causes compositional evolution of freezing solutions. The pathway (1) corresponds to precipitation of Ca and Na sulfates, the pathway (2) is referred to formation of Mg and Na sulfates, and the pathway (3) shows precipitation of hydrohalite. Neutral MgSO 4 type solutions may not fully convert to Mgpoor fluids at depth because trapping of SO 2-4 to low-solubility Ca sulfates could be limited by dissolution of Ca minerals. Partial freezing and migration of fluids could separate low-t and/or deep chloride solutions from sulfate-rich salts and brines. Acidic H 2 SO 4 -HCl Brines Depth Neutral MgSO 4 1 Neutral MgSO 4 -rich 2 Neutral Mg-Na chloride Alkaline Na-Ca sulfate-chloride Alkaline Alkaline Na-Ca 1 3 chloride Ca chloride Freezing 71

72 Fig. 14. The supposed formation of Hesperian massive layered sulfate deposits through subsurface remobilization of solutions formed during a widespread acidic chemical weathering in the Noachian epoch. In the Noachian epoch, neutralized alteration solutions and salts accumulated at depth in the vicinity of upper permafrost boundary and some brines percolated within the permafrost. In the Hesperian epoch, the formation of tectonic depressions together with local warming and melting of the permafrost led to subsurface migration of sulfate-rich solutions followed by their freezing and/or evaporation in topographic lows. The right hand side of the scheme illustrates an initial stage of sulfate deposition. The deposition of salts could have occurred concurrently with deposition of airborne silicate dust. In the Valles Marineris trough system, sulfate-rich deposits formed alongside with subsiding of depressions which caused deformation and erosion of initial deposits. 72