Icarus 207 (2010) 659–674

Contents lists available at ScienceDirect

Icarus

journal homepage: www.elsevier .com/locate / icarus

Diagenetic haematite and sulfate assemblages in Valles Marineris

Leah H. Roach a,*, John F. Mustard a, Melissa D. Lane b, Janice L. Bishop c, Scott L. Murchie d

a Dept. Geological Sciences, Brown University, 324 Brook St., Providence, RI 02912, USAb Planetary Science Institute, 1700 E. Fort Lowell Rd., Suite 106, Tucson, AZ 85719, USAc The SETI Institute & NASA-Ames Research Center, Carl Sagan Center, 515 N. Whisman Rd., Mountain View, CA 94043, USAd Applied Physics Laboratory, 11100 Johns Hopkins Rd., Laurel, MD 20723, USA

a r t i c l e i n f o

Article history:Received 8 May 2009Revised 26 October 2009Accepted 20 November 2009Available online 4 February 2010

Keywords:Mars, SurfaceSpectroscopyMineralogy

0019-1035/$ - see front matter � 2009 Elsevier Inc. Adoi:10.1016/j.icarus.2009.11.029

* Corresponding author.E-mail address: leah_roach@brown.edu (L.H. Roac

a b s t r a c t

Previous orbital mapping of crystalline gray haematite, ferric oxides, and sulfates has shown an associa-tion of this mineralogy with light-toned, layered deposits on the floor of Valles Marineris, in chaos ter-rains in the canyon’s outflow channels, and in Meridiani Planum. The exact nature of the relationshipbetween ferric oxides and sulfates within Valles Marineris is uncertain. The Observatoire pour la Miner-alogie, l’Eau, les Glaces et l’Activite (OMEGA) spectrometer initially identified sulfate and ferric oxides inthe layered deposits of Valles Marineris. The Thermal Emission Spectrometer (TES) has also mappedcoarse (gray) haematite in or at the base of these deposits. We use Compact Reconnaissance ImagingSpectrometer for Mars (CRISM) spectra and Context Camera (CTX) and High Resolution Imaging ScienceExperiment (HiRISE) imagery from the Mars Reconnaissance Orbiter (MRO) to explore the mineralogyand morphology of the large layered deposit in central Capri Chasma, part of the Valles Marineris canyonsystem that has large, clear exposures of sulfate and haematite. We find kieserite (MgSO4�H2O) and ferricoxide (often crystalline red haematite) in the lower bedrock exposures and a polyhydrated sulfate with-out ferric oxides in the upper bedrock. This stratigraphy is duplicated in many other basinal chasmata,suggesting a common genesis. We propose the haematite and monohydrated sulfate formed by diage-netic alteration of a sulfate-rich sedimentary deposit, where the upper polyhydrated sulfate-rich, haema-tite-poor layers either were not buried sufficiently to convert to a monohydrated sulfate or were part of alater depositional phase. Based on the similarities between the Valles Marineris assemblages and the sul-fate and haematite-rich deposits of Meridiani Planum, we hypothesize a common evaporite and diage-netic formation process for the Meridiani Planum sediments and the sulfate-bearing basinal InteriorLayered Deposits.

� 2009 Elsevier Inc. All rights reserved.

1. Introduction There is a need to make a strong connection between thermal

The discovery of coarsely crystalline haematite (Christensenet al., 2000) and sulfates (Gendrin et al., 2005b) in the equatorialregions from martian orbital data from Mars (e.g. Christensenet al., 2001b; Bibring et al., 2006) provided direct evidence of liquidwater having been present on the surface (Christensen et al., 2000,2001a). The in situ investigation by the Mars Exploration Rover(MER) Opportunity discovered significant jarosite and other sul-fates with haematite and confirmed the aqueous origin of theassemblage (Squyres and Knoll, 2005; McLennan et al., 2005). Hae-matite and sulfate have also been mapped from orbit in specificlocations throughout Valles Marineris and the chaos regions be-tween Valles Marineris and Meridiani Planum (Christensen et al.,2001b; Gendrin et al., 2005b; Glotch and Christensen, 2005; Glotchand Rogers, 2007; Weitz et al., 2008; Lichtenberg et al., 2009), sug-gesting that aqueous alteration is widespread.

ll rights reserved.

h).

infrared and VNIR detections of ferric oxide to determine whatphases are present, how they were created, and their geologic con-text. The definitive VNIR identification of the crystalline ferricoxide and its spatial and geologic relationship to the coarsely crys-talline haematite and sulfate deposits are outstanding questions. Isthe ferric phase a surface coating or an in situ alteration product?Are there stratigraphic and genetic relationships between the Val-les Marineris deposits and the haematite and sulfate-rich sedimen-tary sequences in Meridiani Planum?

Here we test the applicability of the Meridiani Planum model tothe ILDs in Valles Marineris. If we can apply the detailed geologicand chemical evidence from the Opportunity rover in MeridianiPlanum toward an understanding of these deposits, our findingswould result in a large step forward in our knowledge of the VallesMarineris geologic history.

We use data from Compact Reconnaissance Imaging Spectrom-eter for Mars (CRISM), which is a VNIR spectrometer aboard MRO,to better resolve the variety of ferric mineralogy present in VallesMarineris. Geologic and mineralogic mapping of all of the ILDs

660 L.H. Roach et al. / Icarus 207 (2010) 659–674

reveal a regionally consistent, repeatable stratigraphy: monohy-drated sulfate and ferric oxide (commonly red crystalline haema-tite) are found in the lower bedrock layers, while polyhydratedsulfate without ferric oxides are found in the upper bedrock layers.A detailed spectral and stratigraphic study in Capri Chasma is pre-sented as a type locality, as Capri Chasma has strong spectral sig-natures and abundant exposed bedrock. The geologicimplications for this stratigraphy in Valles Marineris are discussed,and considered in the context of the Meridiani Planum models asboth surface and orbital data are available there.

2. Background

The Mars Global Surveyor Thermal Emission Spectrometer (TES)identified coarsely crystalline haematite with a surface abundancein Meridiani Planum of 25–60%, which likely required significantaqueous activity (Christensen et al., 2000). Crystalline haematitewas also mapped within Valles Marineris (Christensen et al.,2001b) and in several of the chaos terrains (Glotch and Christen-sen, 2005; Glotch and Rogers, 2007; Noe Dobrea et al., 2008; Weitzet al., 2008).

MER Opportunity landed in Meridiani Planum and confirmedthe coarse haematite detection and an aqueous formation process(Squyres and Knoll, 2005; McLennan et al., 2005). There are twoleading hypotheses for the formation of the haematite and sul-fate-rich stratigraphy in Meridiani Planum. In the first theory, aeo-lian deposition with cementation by a fluctuating groundwatertable (Squyres and Knoll, 2005; Clark et al., 2005; Tosca et al.,2005; McLennan et al., 2005). Under the first hypothesis, iron oxi-des would have formed by dissolution of Fe sulfates (like jarosite)by surface or groundwater with changing geochemistry (Bighamand Nordstrom, 2000; Clark et al., 2005; McLennan et al., 2005;Barron et al., 2006; Golden et al., 2008). The second hypothesis ispyrite-rich pyroclastic ash flows and air fall where haematite formsthrough hydrothermal precipitation (Hynek et al., 2002; McCollomand Hynek, 2005). Both models are consistent with the regionalgeology (Andrews-Hanna et al., 2007; Hynek and Phillips, 2008),and more importantly for this study, require aqueous alterationto form the haematite and sulfates.

Ferric minerals were identified previously in basinal chasmatathroughout Valles Marineris with data from the instrument Obser-vatoire pour la Mineralogie, l’Eau, les Glaces et l’Activite (OMEGA)on board the European Space Agency’s Mars Express spacecraft, butcould not be definitively resolved as due to ferric sulfates, such ascopiapite or schwertmannite, or as a variety of ferric oxides (Bibr-ing et al., 2007). One of the challenges of interpreting sulfate andferric oxide stratigraphy from OMEGA data is that the spectralidentifications are on too coarse a scale to tie to specific geologicunits identified in high resolution imagery (such as Mars OrbiterCamera (MOC) or High Resolution Stereo Camera (HRSC) data).The 400 m to 4 km-sized footprint of the OMEGA instrument couldcover an entire, mineralogically complex outcrop with a handful ofpixels. The OMEGA team used multiple spectral parameters tocharacterize the ferric absorption in the visible-near infrared(VNIR), because spectral matching to library mineral spectraproved inconclusive (Poulet et al., 2007, 2008). These parametersare indicative of mineralogy but require further spectral analysisto confirm the identifications. When these parameters were cou-pled with Modified Gaussian Models, linear mixing, and nonlinearmixing models, the ferric minerals in Valles Marineris and Merid-iani Planum were interpreted as ferric oxides (Bibring et al.,2007; Le Deit et al., 2008; Poulet et al., 2008).

Hydrated sulfates, including monohydrated Mg sulfate (kiese-rite) and a polyhydrated sulfate of unknown cation composition,were detected in light toned deposits throughout Valles Marineris

(Gendrin et al., 2005b). The light toned, layered deposits were clas-sified as Interior Layered Deposits (ILDs) (Lucchitta et al., 1992). Ageneral correlation of ferric oxide with sulfates was noted (Gendrinet al., 2005a; Hutchison et al., 2005; Bibring et al., 2007; Le Deitet al., 2008; Poulet et al., 2008; Mangold et al., 2008), but detailedstratigraphic mapping needed higher resolution spectral data. Fer-ric oxides were often found downslope or at lower elevations thanhydrated sulfates, but sometimes the extent of hydrated sulfatesand ferric oxides overlapped on the slope of the outcrop. In thosecases, it was unclear if the spectral signature represented a hy-drated ferric sulfate (or sulfate assemblage including kieseriteand a ferric sulfate) or kieserite plus a ferric oxide. Where the ferricsignature was separated from sulfate signatures, as at the base ofILDs, it was interpreted as a ferric oxide (e.g. Mangold et al., 2008).

There have been numerous telescopic, orbital, and landed spec-tral studies of the globally homogeneous dust on Mars (e.g. McCordet al., 1977; Pinet and Chevrel, 1990; Mustard and Bell, 1994; Bellet al., 2000; Morris et al., 1997, 2000; Bandfield and Smith, 2003;Yen et al., 2005). VNIR data suggest that the dust is spectrally dom-inated by a nanophase ferric oxide component and has minor con-tributions from crystalline ferric oxide and ferrous phases (Morriset al., 2006; Yen et al., 2005) and possible carbonate (Bandfieldet al., 2003) or iron sulfate (Lane et al., 2002) as suggested bymid-infrared TES data. This dust is present both as a ubiquitouscoating on the surface and as an aerosol particulate in the atmo-sphere (McCord et al., 1977; Wolff et al., 1997). Understandingthe VNIR spectral signature of martian dust is crucial to being ableto resolve other ferric oxide phases, like crystalline haematite, onthe surface.

2.1. Datasets

In this study, we used multiple orbital imagers and VNIR spec-trometers to study the mineralogy and morphology of these depos-its. The CRISM instrument on the MRO spacecraft is a VNIRhyperspectral imaging spectrometer capable of acquiring observa-tions in both targeted and mapping mode. Targeted observationswere taken in 544 wavelengths from 0.392 to 3.92 lm at �20 m/pixel by two independent spectrometers (Murchie et al., 2007).The ‘‘S” detector covered 0.392–1.02 lm in the VNIR while the‘‘L” detector covered 1.02–3.92 lm in the near-infrared (NIR) (Mur-chie et al., 2007). Observations at 72 wavelengths and 100–200 m/pixel were acquired in CRISM’s multispectral survey mode (Mur-chie et al., 2007). The multispectral wavelengths were chosen asa subset of the targeted observation wavelengths to cover impor-tant absorption bands. VNIR spectroscopy is ideal for detectingminerals that have electronic transition absorptions from ironand vibrational overtones and combination tones from anions suchas OH�; SO2�

4 , and CO2�3 (Burns, 1993; Clark et al., 1990). CRISM’s

higher spatial resolution than OMEGA’s allows for the collectionof spectra from distinct units within an outcrop, thus spectra haveless spatial mixing. The higher spatial resolution also allows selec-tion of just the strongest spectral signatures within unit, for morerobust spectral identification.

Two high resolution cameras are coaligned with CRISM on MROand can take coordinated observations with the spectrometer. HiR-ISE acquires imagery at 25 cm/pixel in strips �5 km across (McE-wen et al., 2005). HiRISE data are useful when examining finelayering or detailed stratigraphic relationships. CTX capturesimages at �6 m/pixel and �30 km across (Malin et al., 2007). Mul-tiple CTX scenes can be mosaicked together to create high resolu-tion base maps useful for correlating unit morphologies across longdistances.

The OMEGA hyperspectral spectrometer covers a similar wave-length range to CRISM (352 spectral channels between 0.35 and5.1 lm) with broader spatial coverage. The spatial sampling varies

Fig. 1. Regional unit map of sulfate distribution in Capri Chasma. Kieserite (green),PHS (blue) and mixed mono- and polyhydrated sulfate (yellow). CRISM targetedobservations containing ferric phases are shown in white while those without ferricphases are black. Location of cross-section (Fig. 8) is drawn in black. Geologiccontacts are determined by CRISM spectral analysis and extended over areas ofmorphologic similarity and along areas with similar elevation. (For interpretation ofthe references to color in this figure legend, the reader is referred to the web versionof this article.)

L.H. Roach et al. / Icarus 207 (2010) 659–674 661

from 300 m (at pericenter) to 4.8 km (at 4000 km altitude) (Bibringet al., 2005). OMEGA data are useful in regions without CRISM data.OMEGA data are calibrated to reflectance using the standard steps(as described in Bibring et al. (2005) and Mustard et al. (2005)).

In addition to the imagers and spectrometers, we used topogra-phy data from the Mars Orbiter Laser Altimeter (MOLA) (Zuberet al., 1992) to create interpretive stratigraphic cross-sections.

Fig. 2. Example textures of units within ILDs containing kieserite (top

2.2. Study locale

Capri Chasma, in eastern Valles Marineris, consists of a largelayered sedimentary mound in the center of the canyon sur-rounded by knobs and other collapse features (Fig. 1). The moundis classified as an ILD (Lucchitta et al., 1992). The mound in CapriChasma (300 km across, up to 4 km tall) shows a bright-toned,massive unit at the lower elevations capped by a dustier, slightlylower albedo, finer textured unit above. The massive lower unitshows erosional fluting on exposed faces, while much of the upperunit has a blocky texture or has been eroded away. The lower albe-do of the upper unit may be due to intrinsic properties or may bean effect of the greater dust cover on the generally flatter upperunit. A �15 km wide mound to the northeast may be an aeolianredeposition of the upper unit, as it differs in morphology andhas a lower thermal inertia (Putzig and Mellon, 2007) than surfaceswith a similar mineralogy. The ILD has been suggested to form byaeolian (Lucchitta et al., 1992; Peterson, 1981; Nedell et al., 1987),lacustrine (McCauley, 1978; Lucchitta, 1982; Nedell et al., 1987),sub-ice volcanic (Chapman and Tanaka, 2001; Komatsu and Lita-sov, 2002; Komatsu et al., 2004), or volcanic processes (Lucchitta,1981; Witbeck et al., 1991; Lucchitta et al., 1992; Weitz, 1999).

This ILD is representative of other ILDs located in the centers ofthe broad, basinal chasmata (Fig. 2). Capri Chasma was chosen fordetailed mapping of ferric oxides and sulfates because TES, OME-GA, and CRISM data show strong spectral signatures indicative ofgood bedrock exposure over a large area. More regional (coarserscale) mapping of ferric oxide and sulfate distribution was per-formed throughout Valles Marineris.

row) and PHS (bottom row). All HiRISE images are 1 km across.

662 L.H. Roach et al. / Icarus 207 (2010) 659–674

3. Methods

3.1. CRISM data reduction

CRISM data are reduced to atmospherically corrected apparentsurface (I/F) reflectance via a standard calibration pipeline (Murchieet al., 2007, 2009b). The potential presence of aerosols in the atmo-spherically corrected CRISM data is most important at the shortestwavelengths (0.4–0.6 lm). Caution must be employed in visiblewavelength spectral interpretation, as aerosols dampen surfacereflectance features and may add their own dust-like absorptions.

Fig. 3. RELAB library spectra 0.38–2.5 lm. Ferric oxides on top and Fe and Mgsulfates on bottom. Szomolnokite (BKR1JB622A), jarosite (C1092F53), kieserite(F1CC15), hexahydrite (LASF57A), haematite (F1CC17B), magnetite (CJB258), lep-idocrocite (892F50), ferrihydrite (C1092F55), goethite (C1GO01). (For interpretationof the references to color in this figure legend, the reader is referred to the webversion of this article.)

3.2. VNIR spectroscopy of sulfates and ferric oxides

Analyses of atmospherically corrected I/F spectra and ratios ofthose spectra over the whole VNIR wavelength range are necessaryto determine the specific mineralogy present in a CRISM observa-tion. Spectra are averaged over tens or hundreds of pixels to im-prove the signal to noise. A spectrum of interest is also dividedby a spectrum of a spectrally neutral (i.e. dusty) region to accentu-ate mineral absorption bands and to remove instrumental artifactsand residual calibration effects. The spectrally neutral denominatorspectrum should come from the same column(s) as the spectrumof interest to account for column-dependent variations in thedetector’s wavelength array (Murchie et al., 2007, 2009a). A spec-trally neutral denominator is best, as mineral absorptions in thedenominator show up as inverse features in ratioed spectra andcan complicate mineral identification. In practice, the denominatoris from dusty terrain, which can convolve dust-related spectral fea-tures in the resulting ratio spectra.

The atmospherically corrected I/F and ratioed spectra are thencompared to a spectral library resampled to CRISM wavelengths.The VNIR and IR spectra are analyzed independently. This is be-cause the two detectors’ radiances do not align at the 1 lm join,possibly because of uncorrected effects of the beamsplitter tem-perature (Murchie et al., 2007). Thus, it is not possible to calculatea reasonable continuum across the full wavelength region. Moredetails on the continuum removal process are given after the dis-cussion of relevant spectral features.

The two spectral classes discussed in this paper are hydratedsulfates and ferric oxides. All published library spectra of thesephases were used for mineral identification. Hydrated sulfateshave an absorption near 2.4 lm due to H2O and OH combinationsand sulfate bending overtones (Fig. 3) (Cloutis et al., 2006). Mono-hydrated sulfates have distinct spectral features from those ofpolyhydrated sulfates. Kieserite, MgSO4�H2O and szomolnokite,Fe2+SO4�H2O, are structurally similar and thus have similar combi-nation H2O stretch and rotation vibrations at �2.1 lm. This band israrely observed in mineral spectra and is a clear indicator of mon-ohydrated sulfates. Kieserite and szomolnokite can be distin-guished from each other in natural samples because this band iscentered at 2.12–2.13 lm in kieserite and 2.09–2.10 lm inszomolnokite. Szomolnokite has an additional broad absorptionnear 0.9 lm due to Fe2+ crystal field transitions (Cloutis et al.,2006). On the other hand, polyhydrated sulfates, such as hexahy-drite (Fig. 3) and bloedite (not shown), have strong H2O vibrationsat 1.45 and 1.95 lm, as well as a drop off in reflectance near2.45 lm attributed to S–O or OH/H2O combinations and/or over-tones (Cloutis et al., 2006). The hydrated Ca-sulfates, gypsum (Ca-SO4�2H2O) and bassanite (CaSO4�0.5H2O) (neither are shown), haveadditional features near 2.2–2.25 lm (Cloutis et al., 2006) that dis-tinguish them from the generic polyhydrated sulfates discussedabove.

Ferric oxides have a variety of very distinctive absorption fea-tures due to the iron in their structures. Ferric oxides such as hae-

matite, goethite, magnetite, lepidocrocite, and ferrihydrite havestrong iron crystal field transitions in the 0.9 lm region (Fig. 3)(Burns, 1993; Morris et al., 1985, 1997). Two Fe sulfates,szomolnokite and jarosite (KFe3(SO4)2(OH)6), are also includedfor comparison. Iron oxides owe their red color to saturated UVcharge transfer absorptions that cause very low reflectivity short-ward of 0.55 lm (Morris et al., 1985). Iron oxides show a varietyof �0.5 lm absorptions, which are caused by ferric ion chargetransfers or crystal field splitting (Hunt and Salisbury, 1970; Huntet al., 1971; Morris et al., 1985). Fig. 4 shows RELAB library spectraof ferric oxides and sulfates after continuum removal. Crystallinered haematite (thick red line in Fig. 3A) is distinguishable fromall other ferric oxides and sulfates since it has a 0.53 lm minimumat distinctly longer wavelengths than other minerals. The presenceof two minima (0.53 and 0.88 lm), a 0.62 lm shoulder, and0.74 lm relative peak also provide many features in the visiblewavelengths to fit to and thus the possibility of a confident spectralmatch.

Identification of iron oxide or sulfate species is attempted aftercontinuum removal. The continuum is removed from CRISM visiblespectra over the 0.41–0.97 lm region to better resolve ferricabsorptions in the presence of martian dust. Globally homoge-neous dust on Mars also has a nanophase ferric oxide componentin addition to minor crystalline ferric oxide and ferrous phases(McCord et al., 1977; Pinet and Chevrel, 1990; Bell et al., 2000;Bandfield and Smith, 2003; Lane et al., 2004; Yen et al., 2005). Mar-tian dust is characterized by a broad, shallow ferric absorption cen-tered near 0.9 lm, a local maximum near 0.75 lm, and anabsorption edge shortward of �0.5 lm (Bell et al., 2000; Morris

Fig. 4. Continuum-removed RELAB library spectra 0.41–0.97 lm. (A) Ferric oxides,with lines for 0.53 and 0.88 lm haematite minima. (B) Fe and Mg sulfates. (Forinterpretation of the references to color in this figure legend, the reader is referredto the web version of this article.)

L.H. Roach et al. / Icarus 207 (2010) 659–674 663

et al., 2000; Yen et al., 2005). Since dusty regions in CRISM imagesare used in spectral ratioing to reduce instrumental noise, theabsorptions due to iron in the dust need to be distinguished fromcrystalline ferric phases in the spectrum of interest.

The continuum removal computation fits a convex hull over thespectrum using straight segments connecting local spectral max-ima (ITT, 2007). The first and last value in the wavelength regionundergoing continuum removal is set to 1.0. Two different con-tinua are computed over the VNIR and IR spectral ranges. Wedetermine ferric oxide mineralogy in CRISM ratioed and contin-uum-removed ratioed spectra over 0.41–0.97 lm. We categorizehydrated sulfates based on absorption in CRISM ratioed and con-tinuum-removed ratioed spectra over 1.01–2.5 lm. The CRISMshort wavelength channels between 0.644 and 0.684 lm, less than0.41 lm, and greater than 0.97 lm are routinely excluded fromanalysis. The wavelengths <0.41 lm are degraded due to artifactsfrom the scattered light correction, and the region between 0.644and 0.684 lm has a detector artifact (Murchie et al., 2009b). How-ever, the remaining channels are sufficient to resolve the haematiteand other ferric oxide absorption features. In the IR spectra, thecontinuum is fit over 1–2.5 lm because that is the wavelength re-gion of interest for hydrated mineralogy.

3.3. Mapping methodology

The regional unit map in Capri Chasma was created as follows.First, detailed spectral analysis of CRISM targeted and multispec-tral mapping data and OMEGA data identified locations with kiese-rite and/or PHS. Then HiRISE and CTX data were examined todetermine the association of sulfate mineralogy to morphologicunits. MOLA shot and gridded data were used to determine eleva-

tions of the sulfate-bearing layers. The kieserite and PHS unitswere extended from the CRISM and OMEGA identifications to re-gions with similar morphology and topography. In regions withincomplete CRISM or OMEGA coverage and/or obscuration by adusty, spectrally neutral mantle, the sulfate units were mappedwith the morphology and elevation constraints. CRISM targeteddata in 10 locations and MOLA shot data on both the northernand southern part of the ILD were used to determine the elevationof the contact between kieserite and PHS.

The spectral signatures of haematite or ferric phases cannot betied to a unique morphologic unit in HiRISE or CTX data, nor arethey widespread in CRISM imagery. Thus, ferric phases are notmapped as continuous units; rather, CRISM targeted observationsthat show the presence of red haematite, probable red haematite,or ferric phases are outlined in white, while other CRISM observa-tions lacking ferric spectral signatures are shown in black in Fig. 1.The details of the relationship between sulfates and ferric phasesare best seen in individual CRISM observations.

4. Spectral results

4.1. Minerals identified in Capri Chasma

CRISM type spectra of each phase were chosen over dust-freeareas with the strongest spectral features (see Table 1). Kieserite isidentified in Capri Chasma based upon characteristic H2O vibrationalabsorption features at 2.13 and 2.4 lm (dark gray line in Fig. 5A). TheCRISM spectrum lacks the broad 1.6 lm in the library kieserite spec-trum (dark blue in Fig. 5A). However, the broad 1.6 lm absorption isnot visible in all terrestrial kieserite samples or in all martian spectra.The continuum-removed CRISM spectrum is more consistent withkieserite (2.13 lm minimum) than szomolnokite (2.09 lm)(Fig. 5B). The VNIR region of the CRISM kieserite type spectrum,the red spectrum in Fig. 5, also does not have szomolnokite features.Polyhydrated sulfates (PHS), such as hexahydrite and bloedite, areidentified by absorptions near 1.4 and 1.9 lm due to H2O (Fig. 5).Spectral similarities between PHS preclude identifying a specificphase. The PHS are probably iron-poor, since they lack iron elec-tronic transitions near 0.9 lm. The absorption near 2.24 lm maybe due to admixture with hydrated silica.

A variety of ferric absorptions occur in materials associatedwith the sulfate-bearing ILDs (Fig. 6A). We classify the type spectraas ‘‘red haematite”, ‘‘probable (red) haematite”, ‘‘ferric phases”, and‘‘dust” based on their features in CRISM continuum removed spec-tra (Fig. 6B). Well-crystalline red (i.e., fine-grained, Lane et al.,1999) haematite can be resolved from other ferric oxides and fromferric sulfates by absorptions at 0.53 and 0.88 lm, and a peak at0.76 lm with a shoulder near 0.62 lm (Morris et al., 1985). Contin-uum-removed ratioed CRISM spectra that only have three of thefour haematite features in the right position or all four featuresbut at slightly shifted wavelengths (within 0.02 lm) are classifiedas ‘‘probable (red) haematite”. Spectra with visible wavelengthabsorptions different from the nanophase ferric oxide in the dustare classified as ‘‘ferric phases”. These spectra could be variablemixtures of haematite, ferric sulfates, and dust. The ‘‘dust” spectralclass in Fig. 6B is an unratioed continuum-removed spectrum.‘‘Dust” is not shown as a ratioed spectrum, since similar spectraare used as denominators for the other spectral types. Overall itsspectral features match those reported from previous studies(e.g. Bell et al., 2000; Morris et al., 2000; Yen et al., 2005).

The ability to classify a spectrum as ‘‘red haematite” or as‘‘probable (red) haematite” is predicated on the image’s data qual-ity. A CRISM observation with high quality spectra and a largespectrally neutral (dusty) region in the same columns as the spec-tra of interest is much more likely to resolve all four haematite

Table 1Location of CRISM type spectra.

Mineral typespectrum

CRISMobservation

Detector location

Kieserite FRT000066F5 L detector, 5 � 5 pixel average centered at (68, 258) divided by 5 � 5 average at (68, 39)PHS FRT0000C564 L detector, 5 � 5 pixel average centered at (150, 334) divided by 5 � 5 average at (150, 391)Haematite FRT000066F5 S detector, 5 � 5 pixel average centered at (64, 258) divided by 5 � 5 average at (64, 39). Coaligned with ‘‘kieserite” type

spectrum in L detectorProbable haematite HRS000044E6 S detector, 5 � 5 pixel average centered at (30, 62) divided by 5 � 5 average at (30, 120)Ferric phase FRT0000B385 S detector, 15 � 15 pixel average centered at (149, 378) divided by 15 � 15 average at (149, 157)Dust FRT00009DF9 S detector, 15 � 15 pixel average centered at (40, 32)

Fig. 5. CRISM ratioed (A) and continuum-removed ratioed (B) spectra of typical kieserite/haematite location (gray) and of haematite-poor PHS (black) location in CapriChasma. The kieserite/haematite spectrum has IR absorptions that match the RELAB library natural kieserite spectrum (dark blue) while VNIR absorptions are characteristic ofhaematite (red spectrum in Fig. 4) rather than szomolnokite (light blue in Fig. 4). PHS is an approximate fit to RELAB library hexahydrite (green). (For interpretation of thereferences to color in this figure legend, the reader is referred to the web version of this article.)

Fig. 6. (A) Ratioed CRISM spectra of typical ‘‘red haematite”, ‘‘probable red haematite”, and ‘‘ferric phases” over 0.4–1.1 lm. Denominator for ratioing in all is similar to ‘‘dust”spectral type. (B) Continuum-removed ratioed CRISM spectra of ‘‘red haematite”, ‘‘probable red haematite”, ‘‘ferric phases”, and ‘‘dust” spectral types.

664 L.H. Roach et al. / Icarus 207 (2010) 659–674

spectral features than one that has high aerosol loading, smallexposures of the spectra of interest, and/or no suitable spectrallyneutral region for ratioing. Similarly, further imaging of a locationclassified as containing ‘‘ferric phases” may reveal the spectra areactually consistent with ‘‘red haematite” or ‘‘probable haematite”.Thus, this study is a conservative estimation of the locations ofhaematite deposits in CRISM data.

Ferrous minerals like olivine and pyroxene can be excludedbased on the position and width of the 1.0 and 2.0 lm absorptions.Although these mafic minerals are a dominant phase in the chasma

walls and in volcanic plains across Mars, they are not present inmeasurable amounts in ILDs.

The above criteria for sulfate and ferric oxide spectral identifica-tions were also applied to CRISM spectra across Valles Marineris(including Tithonium, Ius, Melas, Candor, Ophir, Coprates, Ganges,Juventae, and Hebes Chasmata). Table 2 lists 134 CRISM observa-tions in Valles Marineris that show sulfate or ferric oxide features;a total of 299 CRISM observations were examined as part of thisstudy. Duplicate observations were included if they were acquiredunder periods of more favorable aerosol loading and thus had more

Table 2Location within Valles Marineris, the type of sulfate and ferric phase detected, and the geologic setting of the ferric phase within the CRISM observation. The ‘‘sulfate” category isclassified as �1: long wavelength imagery not acquired, 0: no sulfate detected, 1: PHS, 2: kieserite, 3: PHS and kieserite, and 4: sulfate and another hydrated phase. The ‘‘ferricphase” category is classified as �1: short wavelength imagery not acquired, 0: no ferric phase detected, 1: red crystalline haematite, 2: probable haematite, 3: ferric phase, and 4:multiple ferric phases present. Multiple ferric phases could include red haematite plus another deposit with a different ferric absorption, perhaps due to another ferric oxide orsulfate. The ‘‘location of ferric phases” was classified as 1: in bedrock only, 2: in loose material only, or 3: in both bedrock and loose material. In observations where multiple ferricphases were detected, often one type would be in bedrock and the other in loose material.

Latitude Longitude Observation Location Sulfate Ferric phase Location of ferric phase

�12.66 �47.28 FRT0000282F Eos 1 1 1�12.6 �46.68 FRT00003050 Capri 3 3 1�12.88 �48.41 FRT000035BD Capri 3 1 1�4.89 �89.24 FRT00003EEF Tithonium 3 1 3�5.7 �75.6 FRT00004694 Candor 3 2 1�13.1 �64.7 FRT0000474F Coprates 4 0�5.29 �89.24 FRT0000510D Tithonium 3 0�5.79 �75.83 FRT00005350 Candor 1 0�5.26 �76.53 FRT00005521 Candor 3 4 3�4.53 �62.13 FRT00005633 Juventae 3 1 1�4.6 �63.1 FRT00005C2B Juventae 2 �1�5.79 �75.16 FRT00005D17 Candor 1 1 1�11.83 �69.18 FRT00005D59 Melas 1 3 2�10.26 �73.76 FRT000061BD Melas 3 1 1�12.17 �68.94 FRT000061F3 Melas 3 0�6.11 �73.91 FRT00006470 Candor 3 4 3�8.58 �65.54 FRT00006622 Candor 1 0�13.39 �48.66 FRT000066F5 Capri 3 1 1�4.9 �90.26 FRT0000679F Tithonium 3 0�12.53 �46.16 FRT0000681A Capri 1 1 1�2.06 �17.59 FRT00007E16 Iani 1 0�4.64 �89.57 FRT00007EC9 Tithonium 3 0�15.86 �50 FRT00007F63 Eos 3 0�12.51 �46.19 FRT00007FAA Capri 3 1 1�4.68 �63.2 FRT0000803B Juventae 2 2 1�8.26 �66.87 FRT000081A4 Candor 1 0�4.92 �89.68 FRT0000835F Tithonium 3 3 3�6.21 �69.58 FRT0000849B Candor 3 0�12.89 �48.39 FRT000085F9 Capri 2 1 1�7.39 �66.66 FRT0000869B Candor 1 0�6.44 �67.88 FRT00008849 Candor 0 2 3�8.72 �79.21 FRT00008950 Ius 4 0�10.19 �69.34 FRT00008A02 Coprates 2 0�4.65 �26.94 FRT00008C71 Aureum 3 1 3�14.15 �48.06 FRT00008CC0 Capri 2 1 1�9.2 �78.05 FRT00008E66 Ius 1 0�5.36 �89.76 FRT00009390 Tithonium 1 2 1�8.66 �79.24 FRT00009445 Ius 3 0�10.44 �98.47 FRT000096EE Noctis 4 4 3�3.95 �61.88 FRT000097A8 Juventae 2 1 1�5.37 �89.41 FRT000097B1 Tithonium 3 1 3�6.45 �73.68 FRT000098BB Candor 2 1 3�6.26 �69.15 FRT00009901 Candor 3 1 1�6.75 �48.97 FRT00009A1B Ganges 1 2 1�5.81 �76.02 FRT00009A20 Candor 3 4 3�6.35 �70.49 FRT00009A67 Candor 3 2 3�3.95 �61.88 FRT00009ADC Juventae 2 1 1�4.42 �62.31 FRT00009C0A Juventae 3 4 3�5.85 �73.89 FRT00009CB6 Candor 2 4 3�13.89 �58.95 FRT00009D64 Coprates 4 0�4.32 �27.04 FRT00009DAD Aureum 3 3 1�5.81 �76.02 FRT00009DF9 Candor 3 3 3�4.37 �26.8 FRT0000A0B2 Aureum 1 0�7.24 �49.44 FRT0000A0F0 Ganges 3 2 1�9.77 �70.13 FRT0000A131 Melas 4 0�3.92 �26.61 FRT0000A38B Aureum 3 3 1�11.12 �75.11 FRT0000A3E9 Melas 1 0�5.56 �73.48 FRT0000A5B6 Candor �1 2 1�12.88 �49.55 FRT0000A82E Eos 3 3 1�4.69 �72.91 FRT0000A86A Ophir 3 4 3�8.56 �79.46 FRT0000A91C Ius 3 0�6.96 �75.1 FRT0000A938 Candor 1 0�12.35 �69.59 FRT0000A95B Melas 3 3 2�8.91 �77.19 FRT0000AA51 Ius 1 0�6.1 �76.13 FRT0000B0F2 Candor 3 0�4.08 �74.27 FRT0000B27B Ophir 3 0�13.22 �47.46 FRT0000B385 Eos 3 4 3�6.07 �71.05 FRT0000B3C0 Candor 2 0�5.54 �73.25 FRT0000B4D2 Candor 2 2 1

(continued on next page)

L.H. Roach et al. / Icarus 207 (2010) 659–674 665

Table 2 (continued)

Latitude Longitude Observation Location Sulfate Ferric phase Location of ferric phase

�10.36 �68.81 FRT0000B64A Coprates 3 4 3�12.61 �47.2 FRT0000B776 Eos 3 1 1�12.77 �47.5 FRT0000BB25 Capri 3 1 1�5.34 �75.34 FRT0000BB2A Candor 1 2 1�4.37 �71.18 FRT0000BB63 Ophir 2 1 2�5.21 �72.77 FRT0000BD02 Ophir 2 4 3�6.27 �69.36 FRT0000BE68 Candor 3 4 3�12.77 �47.5 FRT0000C152 Capri 3 1 1�6.27 �72.92 FRT0000C433 Ophir 0 2 2�6.28 �69.21 FRT0000C470 Candor 3 4 3�13.06 �47.46 FRT0000C564 Capri 3 2 3�14.3 �48.56 FRT0000C678 Capri 3 1 1�4.35 �62.33 FRT0000C72A Juventae 3 4 3�13.24 �45.55 FRT0000C81D Eos 2 1 1�6.26 �69.38 FRT0000C87C Candor 3 4 3�6.26 �69.36 FRT0000CCB5 Candor 3 4 3�5.19 �89.02 FRT0000CE25 Tithonium 3 1 1�13.39 �48.86 FRT0000D117 Capri 2 1 1�5.99 �73.74 FRT0000D272 Candor 3 1 3�13.33 �47.53 FRT0000D3A4 Capri 3 4 3�8.55 �79.92 FRT0000D5F8 Ius 1 0�12.82 �46.51 FRT0000D62D Capri 2 1 1�5.7 �75.41 HRL00002831 Candor 1 2 1�3.55 �61.75 HRL000028A6 Juventae 2 2 1�6.67 �75.77 HRL000033B7 Candor 1 3 3�10.23 �68.82 HRL00003752 Coprates 2 4 3�3.28 �61.73 HRL0000444C Juventae 2 2 1�4.64 �89.57 HRL00004893 Tithonium 3 3 3�7.81 �64.89 HRL00004F38 Candor 2 4 3�4.25 �71.58 HRL0000508A Ophir 1 0�4.37 �72.61 HRL00005B82 Ophir 2 1 1�3.61 �26.27 HRL00006181 Aureum 3 3 1�7.13 �49.44 HRL0000633F Ganges 2 2 1�7.35 �45.15 HRL0000638C Ganges 1 0�5 �89.74 HRL000063F1 Tithonium 3 3 3�6.85 �68.99 HRL0000648F Candor 3 1 1�6.23 �69.36 HRL00006844 Candor 4 1 3�4.92 �89.68 HRL00007692 Tithonium 3 3 3�4.92 �90.13 HRL00008F8E Tithonium 3 1 2�4.67 �88.81 HRL0000A1BA Tithonium 1 0�13.83 �50.41 HRL0000A23F Capri 3 2 3�4.32 �71.35 HRL0000A432 Ophir 1 3 2�2.51 �61.66 HRL0000A4CF Juventae 3 2 3�7.02 �74.1 HRL0000AB7A Candor 2 1 1�12.01 �46 HRL0000AF21 Capri 0 4 3�4.72 �71.94 HRL0000B7D4 Ophir 2 0�5.14 �73.85 HRL0000C03D Candor 3 1 1�10.23 �69.22 HRL0000C194 Coprates 1 0��8.64 �76.83 HRL0000C2BA Melas 1 0�4.82 �89.97 HRL0000C72F Tithonium 3 3 3�8.13 �79.59 HRL0000C7C6 Ius 2 0�4.82 �89.97 HRL0000CBC8 Tithonium 3 3 3�5.07 �89.3 HRL0000D1BE Tithonium 2 1 3�1.1 �75.41 HRL0000D632 Hebes 3 3 1�13.33 �48.52 HRL0000D770 Capri 3 1 1�12.36 �69.26 HRS0000285D Coprates 3 2 1�12.28 �45.99 HRS00002F87 Capri 1 3 3�6.23 �69.23 HRS00002FAF Candor 3 3 4�10.26 �73.76 HRS000044E6 Melas 3 1 1�5.49 �75.89 FRT0000BE37 Candor 3 3 2�5.71 �75.52 FRT0000C9E8 Candor 2 2 1�5.11 �76.48 FRT0000CD9C Candor 2 4 3�4.89 �89.22 FRT0000ACC1 Tithonium 3 1 3�2.51 �61.64 FRT0000A4CF Juventae 3 2 3�5.29 �75.8 HRL00007C86 Candor 2 3 2

666 L.H. Roach et al. / Icarus 207 (2010) 659–674

clear spectral features. There are 24 observations in Capri/EosChasma that show sulfate or ferric oxide mineralogy (Fig. 1).

4.2. Ferric oxide variability throughout Valles Marineris

Many Interior Layered Deposits (ILDs) in Valles Marineris showabsorptions due to iron electronic transitions, likely crystalline red

haematite and other ferric oxides. A complete list of CRISM tar-geted observations showing ferric oxide is given in Table 2.

Differences in continuum-removed ratioed VNIR spectra aresubtle due to the ubiquitous, muting presence of dust. Diffusedeposits often have weak spectral features where haematiteabsorptions cannot be differentiated from other ferric oxides orferric phases. Thus, concentrated deposits of ferric oxide areneeded to be able to confidently identify them as red haematite.

Fig. 7. (A) Contact between kieserite/haematite and PHS in subset of HiRISEPSP_008088_1760 colorized with CRISM FRT0000A82E. R: shoulder at 0.60 lm, G:1.90 lm band depth, B: 2.10 lm band depth, indicative of haematite, PHS, andkieserite-bearing material respectively. Box shows extent of C. (B) Detail showingdifferent erosional texture of brighter, smoother kieserite and rougher PHS. (C)Closeup of contact between kieserite/haematite (bright, fluted material to NW) andoverlying PHS (darker, more debris-covered material to SE). Box indicates locationof B.

L.H. Roach et al. / Icarus 207 (2010) 659–674 667

One example is in West Candor Chasma, where Murchie et al.(2009a) have shown a concentration of red haematite in thebraided channels eroding into Candor Mensa and a concentrationof ‘‘ferric oxide” at the base of the ILD. This ‘‘ferric oxide” corre-sponds spatially to a TES detection of coarsely crystalline haema-tite. The coarsely crystalline haematite at Meridiani Planum is amixture of gray crystalline haematite abrading to red haematite(Glotch and Kraft, 2008). Gray haematite is spectrally neutral inthe VNIR (Lane et al., 1999), but perhaps the gray haematite depos-it at the base of the ILD is similar to Meridiani Planum deposits and

Fig. 8. Stratigraphic cross-section of sulfates within the central mound in Capri Chasma, fbase of slope by impact processes. Lower extent of kieserite uncertain. (For interpretativersion of this article.)

contains some abraded pieces displaying ferric absorption features.These features may not be sharp enough to resolve as red haema-tite but are spectrally different from dusty spectra used in ratioing.

4.3. Limitations on spectral interpretation

It is common for observations with ‘‘red haematite” or ‘‘proba-ble haematite” to also have other ferric phases present in the scene.The ‘‘red haematite” may be in the bedrock or collecting in chan-nels on the bedrock, while the ‘‘ferric phases” are more commonlyfound in loose deposits at the base of the ILD. The spectra classifiedas ‘‘ferric phases” may contain other ferric oxides, ferric sulfates, ora mixture of ferric phases.

Concentration of the material into channels on the ILD and as alag at the base may have been necessary to get strong enough spec-tral features for identification with CRISM data. We recognize anidentification bias toward ferric phases concentrated into lagsand realize the true iron-bearing mineral diversity in ILDs is likelymuch higher than presented here.

5. Stratigraphic results

5.1. Stratigraphic section and regional unit map of Capri Chasma

The spectral identifications from individual CRISM targeted andmapping observations and OMEGA observations in Capri Chasmaare extended into a regional unit map (Fig. 1). The entire centralILD is interpreted to be sulfate rich.

In Capri Chasma, as elsewhere in the basinal chasmata, the min-eralogic units also exhibit general textural differences. The PHS-bearing units have a rougher texture and thus would trap moreaeolian material than the smoother kieserite-bearing units(Fig. 7B). The smoother texture of kieserite-bearing units may pre-vent aeolian material from collecting so uniformly over the surfaceand may restrict aeolian deposits to local depressions, some ofwhich can be seen in the upper left of Fig. 7C. Aeolian material col-lects along the PHS-rich slope in the center of Fig. 7C. Mangoldet al. (2008) noted that PHS were systematically darker than kiese-rite in West Candor Chasma and similarly suggested either anintrinsic or dust-related cause. The PHS is also often on shallowerslopes that kieserite here and elsewhere in Valles Marineris (Man-gold et al., 2008), which would contribute to its dustiness.

The kieserite-rich material and PHS-rich material show very dif-ferent erosional styles. Large fluted erosional structures cut acrossboth units in the ILD in Fig. 7, but detailed imagery shows that theflute ridge in the kieserite-rich material is smooth and in the PHS-rich material is fractured and breaking into meter-scale blocks(Fig. 7B).

PHS lies stratigraphically above kieserite throughout CapriChasma (Fig. 8). The elevation of the contact between the sulfatesis between �2550 and �850 m in CRISM observations that span

rom MOLA orbit 16830. Kieserite is green and PHS is blue. PHS has been deposited aton of the references to color in this figure legend, the reader is referred to the web

Fig. 9. (Left) False-color CRISM observations FRT0000B776, FRT0000BB25, FRT0000C564, and FRT0000B385 top to bottom (RGB: 2.5, 1.5, 1.08 lm) covering contact betweenkieserite and PHS on the northern edge of Capri Chasma ILDs. Putative 30 km impact crater in southwest of image. (Right) RGB: band depths at 0.86, 1.9, and 2.1 lm, in thiscase indicating haematite, PHS, and kieserite, respectively. Kieserite and haematite-bearing locations (magenta) are common, while PHS and haematite-bearing locations(yellow) are absent. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

668 L.H. Roach et al. / Icarus 207 (2010) 659–674

120 km across the deposit, with systematically lower elevations to-ward the west. Assuming the PHS and kieserite layers are continu-ous across the ILD, that corresponds to a long-range dip of thecontact of <1� westward – essentially flat-lying. MOLA shot datahave 150 m footprints, which is much larger than the contact be-tween the two sulfates. This adds uncertainty to the elevation data,especially as the contact occurs on slopes ranging from 25� to 35�.Fig. 8 shows the slope of ILD flanks and the dip of the contact be-tween kieserite and PHS units. In regions with multiple adjacentCRISM observations that resolve the contact between kieseriteand PHS units, the local dip of the contact is <2� (as in Fig. 9).

The bottom of the kieserite-rich layer is not exposed anywherein Capri Chasma and loose materials or dust mantle the bottom ofILD slopes and surrounding chasma floor. As a result, the totalthickness and lateral extent of the kieserite layer cannot be as-sessed. The maximum thickness of exposed kieserite-rich materialis 2000 m. The PHS-rich unit varies in thickness due to aeolianstripping of the top of the ILD; in some parts of the ILD it is entirelyremoved. The PHS-rich unit has a maximum observable thicknessof 400 m.

5.2. Haematite and sulfate association in Capri Chasma

The northern edge of the central ILD is a key outcrop that clearlyshows the contact between underlying the kieserite- and haema-tite-bearing unit and the overlying PHS-bearing, ferric oxide-poorunit. The stratigraphy of the PHS-bearing unit above the kiese-rite/haematite-bearing unit is consistent throughout the chasmabut is best expressed in CRISM observation FRT0000BB25 andnearby observations (Fig. 9). The northern part of Fig. 9 has anundisturbed stratigraphy while the southern part covers a 30 kmdiameter putative impact crater and has a jumbled mineral stratig-raphy. The sulfate-rich material within the putative crater was pre-viously interpreted to be brecciated bedrock (Roach et al., 2007).The contrast in mineral stratigraphy between the two parts indi-cates the sulfates are a bulk component of the bedrock that isredistributed by impact processes and not due to a surface–atmo-spheric interaction.

Fig. 9B shows band depths indicative of haematite in red(0.86 lm), PHS in green (1.9 lm) and kieserite in blue (2.1 lm).Thus, regions with 0.86 lm and 2.1 lm absorptions (colored ma-genta) contain both kieserite and haematite. Along the northernpart of Fig. 9 where the sulfate stratigraphy is undisturbed, allkieserite-bearing bedrock materials have a ferric oxide component.Spectral analysis shows it is due to red crystalline haematite. Thebase of the kieserite/haematite unit where it meets the canyonfloor does not have either 0.86 or 2.1 lm absorptions. The spectralsignature of the PHS overlying the kieserite lacks significantadmixture with crystalline red haematite – it has a 1.9 lm absorp-tion but not a strong 0.86 lm absorption. This association of hae-matite with kieserite-bearing bedrock units but not with theoverlying PHS-bearing bedrock is seen in CRISM observationsacross Capri Chasma.

Impact cratering can disturb the general stratigraphy of PHSabove kieserite (SW portion of Fig. 9). The cratering and slumpingprocess has reworked the mineral stratigraphy. The 0.86 lmabsorption characteristic of haematite is found alone (red inFig. 9B) and with mono- and polyhydrated sulfates in regions dis-turbed by cratering. Regions that have a 2.1 lm absorption fromkieserite but no 0.86 lm (blue in Fig. 9B) are also common withinthe crater.

Another example of the strong correlation of haematite withkieserite but not PHS is found in the northwest part of the centralILD (Fig. 7). HiRISE image PSP_008088_1760 is colorized withCRISM false color. Similar to Figs. 9B and 7A has a false colorRGB indicating haematite, PHS, and kieserite-bearing units respec-tively. The red channel is the 0.6 lm haematite shoulder ratherthan the 0.86 lm haematite absorption, because in this scene the0.6 lm absorption is a more robust indicator of pixels with thespectral properties of red haematite. However, the red channel alsoshows pixels with a 0.6 lm feature due to the nanophase ferricoxide in the dust. Spectral analysis is necessary to identify pixelswith red haematite or with dust. Thus, the reddish pixels at thebase and on the top of the ILD show spectral features indicativeof dust rather than of red haematite or other ferric phases. Thegreen (1.9 lm) and blue (2.1 lm) channels remain the same as

Fig. 10. Coarsely crystalline haematite distribution in Capri Chasma from TES data. Box indicates extent of Fig 9.

L.H. Roach et al. / Icarus 207 (2010) 659–674 669

Fig. 9B, and indicate PHS and kieserite mineralogy, respectively.The lower part of the ILD has 0.6 and 2.1 lm absorptions from hae-matite and kieserite (colored magenta), while the upper part of theILD has only a 1.9 lm absorption from PHS (colored green). Thelack of red haematite in the PHS units can be seen visually bythe lack of yellow hues.

The haematite is concentrated in the kieserite-bearing bedrockand at the base of some, but not all, kieserite-bearing ILDs. Thisvariation in haematite distribution at the base of ILDs may bedue to differences in aeolian transport of haematite fines to/fromthe base of the ILD. Some locations with haematite in loose mate-rial at the base of an ILD also show a PHS signature in the loosematerial. This may be aeolian deposition of the stripped upper lay-ers plus the haematite-rich lag from erosion of the lower layers.Kieserite is not found in loose material in Capri Chasma.

A map of coarse crystalline haematite as detected by TES(Fig. 10) shows that the phase concentrates at the base and thelower flanks of the northern (well-exposed) wall of the centralILD, but not in the PHS-rich upper unit. Details of how this TES hae-matite index map was constructed can be found in Weitz et al.(2008) in which similar haematite maps were constructed for otherparts of Valles Marineris. Coarse crystalline haematite is also foundin the slumped material in the putative impact crater discussedpreviously. Red (finer-grained) crystalline haematite in CRISM datais located in the bedrock of the central ILD and also sometimes atthe base of the ILD. The spatial correlation between red and coarsehaematite is not exact, but generally, locations with red haematiteoverlap areas with coarse haematite, suggesting the red haematitemay be physically weathered from the coarser haematite. The arealcoverage of coarse haematite is greater, if only because TES hasmore spatial coverage in Capri Chasma than CRISM targetedobservations.

6. Discussion

6.1. Interpreted stratigraphic section across Valles Marineris basins

Regional maps of sulfate distribution in Western and EasternValles Marineris (Fig. 11A and B) from CRISM and OMEGA data fol-low the general distribution shown by previous maps of OMEGAdata (Gendrin et al., 2005b).

Sulfate is concentrated in bright-toned mounds in the basinalchasma but is also found in a few small and low relief floor depos-

its in the linear chasma. In the maps, which use a different colorscheme than the false-color CRISM data, kieserite-rich depositsare colored green, PHS-rich deposits are blue, units with bothkieserite and PHS or with uncertain sulfate mineralogy are yellow(Fig. 11). Bright, light-toned ILDs that appear relatively dust-free inHiRISE imagery but which do not have sulfate signatures in CRISMor OMEGA are colored a transparent gray. These locations may besulfate-free, have dehydrated sulfates, have sulfate in too low con-centrations for detection, or may have enough dust cover to com-pletely mask any spectral signatures from the bedrock. Similarmuting mechanisms were proposed for the weak spectral signa-tures of sulfate in Meridiani Planum from orbital spectrometers(Arvidson et al., 2006). The other possible interpretation for ILDslacking hydration signatures is that they host a spectrally neutral,dehydrated evaporite, like halite or anhydrite. This sulfate-freearea also has been mapped to be free of coarse haematite (Weitzet al., 2008).

The large basinal chasma in Juventae, Ganges, Melas, and WestCandor all have similar sulfate stratigraphy to the central ILD in Ca-pri Chasma. These basinal chasma have bright, mound-like centralILDs 2–4.5 km in thickness. The lower exposures are monohy-drated Mg sulfate- and red crystalline haematite-rich, and theupper exposures contain PHS with no ferric phases. In all chasma,the ferric phase associated with kieserite is red crystalline haema-tite, but in some locations within the chasma, the ferric absorptionis less diagnostic of red haematite and is instead assigned to a moregeneric ‘‘probable haematite” or ‘‘ferric phase” interpretation. Thedegree of layering and the amount of admixture with basaltic sandvaries between the ILDs. However, the erosional style of the basinalILDs are broadly similar – a fluted or channeled morphology withfew boulders on the slopes or at the base, suggesting a fine-grainedcomposition. These ILDs are similar in shape, morphology, locationwithin their respective chasma, and sulfate and haematite stratig-raphy; they may share a common geologic history.

6.2. Implications for formation

There are five key elements to the sulfate stratigraphy: (1) crys-talline red haematite associated with kieserite-bearing bedrock butnot with PHS; (2) PHS on top of kieserite with a horizontal contact;(3) other ferric oxides may be present in loose material at base ofILDs; (4) coarsely crystalline haematite in a lag deposit at the baseof ILDs but sometimes on the ILD flank; and (5) stratigraphy is con-

Fig. 11. Sulfate occurrences in (A) Eastern and (B) Western Valles Marineris in CRISM data. Kieserite-rich deposits are colored green, PHS-rich deposits are blue, units withboth kieserite and PHS or with uncertain sulfate mineralogy are yellow. Bright, light-toned ILDs that appear relatively dust-free but which do not have sulfate signatures arecolored with white dots on a light gray background. Outlines of CRISM observations showing ferric mineralogy in white boxes, and CRISM observations without ferricmineralogy shown in black boxes. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

670 L.H. Roach et al. / Icarus 207 (2010) 659–674

sistent and repeated across 120 km in Capri and in other basinalchasmata. Fig. 12 shows a schematic of how the basinal ILDs couldhave formed and been enriched in sulfates. These elements to-gether suggest that the kieserite and crystalline red and gray hae-matite could have formed by diagenesis of sulfate-bearingsediments (e.g. Tosca et al., 2008) as a consequence of burialand/or a higher heat flow early in the chasma’s history. This diage-netic process would be active in the lower parts of any ILDs thickerthan a couple of kilometer; upper units would remain unaltered.The coarse haematite weathers out of the ILD and collects as alag at the base, while the red crystalline haematite is most concen-trated in the bedrock. The red haematite could be either an ero-sional product of the coarse haematite or also formed bydiagenesis. Both hypotheses are suggested for the red haematiteat Meridiani Planum (Morris et al., 2006). As suggested in (Roachet al., 2009a), there are several possibilities for the formation ofthe overlying PHS: (1) it could have been the original sulfate pre-cipitated which never altered to kieserite, (2) it could have formed

from kieserite by later interaction with near-surface groundwateror snow/water ice or by sulfate dissolution and re-precipitation,or (3) it could be a later precipitate after the kieserite and haema-tite were emplaced.

The difference in erosional style of more competent kieseriteand more friable PHS may be a clue about how the PHS formed –rehydrated kieserite or primary precipitation. Polyhydrated Mgsulfates, such as epsomite (MgSO4�6H2O) or an amorphous, hy-drated Mg sulfate are expected to be the stable phases on Mars,not kieserite (Vaniman and Chipera, 2006). The volume expansionof pure kieserite to crystalline polyhydrated Mg sulfates can resultin expansion by 10–38%. The 35–40% sulfate abundance in Merid-iani Planum (McLennan et al., 2005, Glotch et al., 2006) may be areasonable estimate for the sulfate abundance in ILDs. If so, thencomplete hydration of the kieserite in an ILD could create �10%volume expansion. Hydration of a thin surface layer of the ILD toa polyhydrated Mg sulfate would not cause enough expansion tofracture the surface into 3–4 m-scale blocks, as is seen in Capri

Fig. 12. Schematic of ILD formation and sulfate enrichment in basinal chasmata.Chasma walls not shown. Sulfate enrichment may postdate early ILD formation.

L.H. Roach et al. / Icarus 207 (2010) 659–674 671

Chasma. Thus, if the fracturing is due to rehydration, the PHS in Ca-pri Chasma is present as more than a coating. This PHS is morelikely formed by primary precipitation or dissolution and re-pre-cipitation, rather than atmospheric rehydration. However, the fria-ble nature of PHS-containing units may be due to other factorsthan sulfate hydration state; physical properties such as grain size,packing, and differential stress between layers could also accountfor the difference between the blocky PHS unit and smooth kiese-rite unit.

The strong correlation between kieserite and red haematitesuggest that a common mechanism was responsible for formingthem. Meridiani Planum has a similar association of coarse haema-tite (�10%) and jarosite and other sulfates (35–40%) (Christensenet al., 2004; McLennan et al., 2005; Glotch et al., 2006). The haema-tite concretions are thought to form by shallow groundwater inter-actions in a sulfate-rich basaltic sedimentary sequence (McLennanet al., 2005) or dehydration of goethite to haematite (Glotch et al.,

2004) or alteration of Fe2+ sulfates like melanterite (Fe2+SO4�7H2O)to haematite via dehydroxylation (McLennan et al., 2005; Sefton-Nash and Catling, 2008). In Valles Marineris, low temperature syn-depositional diagenesis of a kilometer-thick sulfate sequence coulddehydrate a polyhydrated Mg sulfate to kieserite, and could trans-form ferric sulfates or oxides to haematite as interstitial cement orconcretions. The ferric oxide component in the global dust may befrom abrasion of Hesperian-aged haematite and not slow surface–atmospheric interactions (Banin et al., 1993).

6.3. Evaporite diagenesis

The general distinction between kieserite at lower elevationsand PHS at higher elevations in basinal chasmata could be due tochanges in climate or groundwater discharge rate or due to post-depositional diagenesis. The sulfates could be an evaporite se-quence with significant aeolian material contribution or an ashand aeolian deposit possibly altered and then cemented by sulfatesthrough groundwater interaction. The source of sulfur could beeither igneous sulfides in the deposit or atmospheric SO2 from vol-canic outgassing (Burns and Fisher, 1990; King and McSween,2005; Chevrier et al., 2006), or both. Changes in aqueous chemistry,such as the freshening of the interstitial brine composition in theupper layers due to meteoritic water input, could partially explainthe enhancement of polyhydrated sulfate at the higher elevationsand monohydrated sulfate at lower. However, in a multi-kilometerthick deposit, syngenetic and diagenetic alteration is unavoidableand can also produce stratified secondary mineralization (Sonnen-feld, 1984).

Once sulfate is precipitated, it is susceptible to dissolution orreaction with brines of many different chemistries (Sonnenfeld,1984). Syngenetic brines derived from the same reservoir that pre-cipitated the sulfates may become more concentrated and sinkthrough the deposit, percolating through unconsolidated evapor-ites and causing mineral recrystallization (Stewart, 1956; Sonnen-feld, 1984). When the deposits are sufficient thickness, lowerportions are lithified by the overburden pressure and increasedtemperatures. This is hypothesized to have occurred in MeridianiPlanum by McLennan et al. (2005). Many of the basinal ILDs onMars are many kilometer thick, and diagenetic processes werelikely active at the lower layers. Brines released by diagenesis areexpelled upward by the lithostatic pressure and can alter overlyingevaporites to secondary precipitates (Sonnenfeld, 1984). After thesulfates became lithified by diagenesis, they are still subject tobrine interactions; epigenetic brines, which are groundwatersunrelated to the evaporite deposit, may percolate through theILD and affect the hydration state or chemistry of the evaporites.

The exact diagenetic temperature or overburden pressureneeded to convert from possible epsomite and other polyhydratedsulfates to kieserite in a multi-sulfate assemblage is complicatedby mineral and groundwater chemistry. One study of a terrestrialseawater-composition brine finds epsomite would convert tokieserite between 25 and 40 �C (Borchert and Muir, 1964). Assum-ing a geothermal gradient of about 10 K/km for the Valles Marine-ris region inclusive of the residual Tharsis volcanic thermal pulse(Schultz and Lin, 2001; Hoffman, 2001), diagenetic conversion tokieserite only requires burial to a few kilometer. Since many ofthe basinal ILDs are several kilometer thick, the kieserite stabilityfield (Borchert and Muir, 1964) may be easily reached in the lowerparts of those ILDs. The increased temperature and pressure at thebase of ILDs would not just have converted more hydrated sulfatesto kieserite and/or haematite, it also would prevent that monohy-drate from becoming polyhydrated by potential late stage flushingwith more diluted brines.

Several of the deposits, such as in Juventae and Melas Chasmata,have very sharp and traceable contacts between kieserite and

672 L.H. Roach et al. / Icarus 207 (2010) 659–674

polyhydrated sulfate. These contacts often have a tonal difference,but not unconformities or other obvious physical disruptions. Thelack of a gradational diagenetic boundary, with co-occurring poly-and monohydrated sulfates, suggests the alteration process wasnot just burial and dewatering, but that groundwater table fluctu-ations may have had a strong imprint of resulting mineralogy.

Other ILDs show alteration between poly- and monohydratedsulfate with stratigraphic section (e.g., eastern and western Can-dor) that cannot be explained by diagenesis alone. The upper layersof Ceti Mensa in Candor Chasma show sharp contacts between twopolyhydrated sulfate layers within a mostly monohydrated sulfate-rich deposit (Murchie et al., 2009a). It is proposed to be due toeither sequential evaporation events or partial diagenesis (Murchieet al., 2009a). The polyhydrated sulfate layers are proposed to beevaporites formed under the wettest conditions. Perhaps the ele-vated temperatures at the bottom of thick ILDs could have pre-vented that monohydrate from becoming polyhydrated duringlate stage flushing with more diluted brines that could depositpolyhydrated sulfate in the upper stratigraphic sections.

6.4. Connections to Meridiani Planum

The layered sulfates at Meridiani Planum have seen similar pro-cesses, with multiple generations of groundwater fluctuationsalternately precipitating, recrystallizing, and dissolving evaporiteminerals and precipitating haematite spherules (Squyres and Knoll2005; McLennan et al., 2005; Squyres et al., 2006). The Whatangacontact of the Burns Formation at Meridiani Planum (�10s cmthick) is interpreted as the terminal height of a water table fluctu-ation and thus as a diagenetic boundary (McLennan et al., 2005;Grotzinger et al., 2005). The Whatanga contact is physically differ-ent than adjacent layers, with a darker albedo and overprinted tex-ture from recrystallization (McLennan et al., 2005; Grotzingeret al., 2005), and chemically different, as represented by variationin Br and Cl concentrations above and below the contact (Clarket al., 2005).

The Meridiani deposits cover 2 � 105 km2 (Hynek, 2004). If theMeridiani Planum model is also applicable to the sulfate-rich Inte-rior Layered Deposits in Valles Marineris, then the region affectedby acidic, aqueous processing is extended by thousands of kilome-ter. Global hydrologic modeling has shown Valles Marineris andassociated chaos to be locations of groundwater upwelling, whichwould have precipitated large volumes of sulfates, oxides, andother phases in solution (Andrews-Hanna et al., 2007). There is agood correlation between predicted upwelling locations and thelarge sulfate deposits in Valles Marineris (Murchie et al., 2009a).The link between Meridiani and Valles Marineris is also supportedby the unusual spectral character of the coarse haematite in bothlocations (i.e., the coarse haematite in both areas lacks a390 cm�1 band Lane et al., 2002; Glotch et al., 2004; Weitz et al.,2008). A common genesis for the Meridiani and Valles Marinerissulfate- and haematite-bearing deposits suggests a broad regionalaqueous environment rich in sulfate with a complex water geo-chemistry that varied with time.

7. Conclusions

Mapping of the mono- and polyhydrated sulfate distribution inILDs across Valles Marineris based on CRISM targeted and mappingobservations shows remarkable consistency in relative mineralogicstratigraphy and trends in mineralogy, deposit texture and ero-sional style. To first order, kieserite is found in layers that arestratigraphically below polyhydrated sulfate-rich layers, especiallyin ILDs greater than several kilometer thick. Red haematite is oftenassociated with kieserite in the ILD bedrock in the CRISM and OME-

GA data (and coarse haematite is associated with the kieserite bed-rock as mapped here and seen in the Mini-TES data) (Christensenet al., 2004), and other ferric oxides are found in either ILD bedrockor in loose material at its base. This mineralogic grouping at thelowest exposed elevations of many ILDs in the broad basins withinValles Marineris (Melas, West Candor, Ganges, Juventae, and HebesChasmata) leads us to hypothesize a regional diagenetic processemplacing haematite and kieserite in previously polyhydrated sul-fate-bearing sedimentary deposits (see Fig. 12).

The ILDs in Valles Marineris are Hesperian-aged sedimentarydeposits that occur over a range of elevations. Taking into accountthe ILD morphology and the common mineral stratigraphy indicat-ing diagenetic alteration of the hydrated phases, we propose amodel for the aqueous geologic history acting on these deposits.They may have originally started accumulating in isolated ances-tral basins and rifts predating the present Valles Marineris config-uration (Schultz, 1998; Lucchitta et al., 1994). Many of the basinalILDs sit unconformably on chaotic terrain. The chaotic terrain maybe the result of collapse from melting of ground ice. That releasedmelt water (or recycling of that water) may be the source of waterfor cementing the ILDs. The ILD deposition and lithification pro-cesses could either have pre-dated or been coincident with originalsulfate enrichment. As the basinal ILDs grew in thickness to a cou-ple kilometers by aeolian, lacustrine, or volcanic processes, diage-netic alteration of sulfate within the lower elevations of thesedimentary structure was unavoidable. The increased tempera-ture (tens of degrees) and overburden pressure would convertpolyhydrated sulfate to kieserite and promote low-temperaturealteration of iron oxides or iron-bearing sulfates such as jarositeto haematite (e.g. Golden et al., 2008). Gradation of the upper por-tions of ILDs have since exhumed the lower, kieserite- and haema-tite-bearing layers. Physical erosion has also concentratedhaematite and ferric oxide as a lag deposit at the base of manyILDs. The PHS-rich units on top of the kieserite-rich units may bethe remnants of upper layers that were not buried sufficiently toconvert to kieserite or they may be part of a later depositionalevent.

This model explains many of the thick ILDs located in the VallesMarineris basins, like Melas, Capri, and Candor Chasmata but can-not explain all sulfate-bearing outcrops; some may be due to dif-ferent processes. For example the character of the �100s m thickILD in Ius Chasma, a linear trough, is consistent with near-com-plete evaporation within a closed basin (Roach et al., 2009b). How-ever, the thick ILDs in the basinal chasma show a repeated mineralstratigraphy and may all be formed by low temperature diagenesis.

We hypothesize that the sulfate-bearing ILDs in Valles Marine-ris and Aram Chaos are genetically related to the sulfate and hae-matite mineralogy in Meridiani Planum and may be the result ofregionally similar aqueous geochemistry and formation processes.

References

Andrews-Hanna, J., Phillips, R., Zuber, M., 2007. Meridiani Planum and the globalhydrology of Mars. Nature 446, 163–166.

Arvidson, R., and 25 colleagues, 2006. Nature and origin of the hematite-bearingplains of Terra Meridiani based on analyses of orbital and Mars ExplorationRover data sets. J. Geophys. Res. 11, E12S08. doi:10.1029/2006JE002728.

Bandfield, J., Smith, M., 2003. Multiple emission angle surface–atmosphereseparations of Thermal Emission Spectrometer data. Icarus 161, 47–65.

Bandfield, J., Glotch, T., Christensen, P., 2003. Spectroscopic identification ofcarbonate minerals in the martian dust. Science 301, 1084–1087.

Banin, A., Ben-Shlomo, T., Margulies, L., Blake, D., Mancinelli, R., Gehring, A., 1993.The nanophase iron mineral(s) in Mars soil. J. Geophys. Res. 98 (E11), 20831–20853.

Barron, V., Torrent, J., Greenwood, J., 2006. Transformation of jarosite to hematite insimulated martian brines. Earth Planet. Sci. Lett. 251, 380–385. doi:10.1016/j/epsl/2006.09.022.

Bell, J., and 23 colleagues, 2000. Mineralogical and compositional properties ofmartian soil and dust: Results from Mars Pathfinder. J. Geophys. Res. 105, 1721–1755.

L.H. Roach et al. / Icarus 207 (2010) 659–674 673

Bibring, J.-P., 10 colleagues, and the OMEGA Team, 2005. Mars surface diversity asrevealed by the OMEGA/Mars Express observations. Science 307, 1576–1581.

Bibring, J.-P., Langevin, Y., Mustard, J., Poulet, F., Arvidson, R., Gendrin, A., Gondet, B.,Mangold, N., Pinet, P., Forget, F., and the OMEGA team, 2006. Globalmineralogical and aqueous Mars history derived from OMEGA/Mars Expressdata. Science 312, 400–404.

Bibring, J.-P.and 11 colleagues, 2007. Mars surface diversity as revealed by theOMEGA/Mars Express observations. Science 317, 1206–1210.

Bigham, J.M., Nordstrom, D.K., 2000. Iron and aluminum hydroxysulfates from acidsulfate waters. In: Sulfate Minerals: Crystallography, Geochemistry, andEnvironmental Significance, Reviews in Mineralogy and Geochemistry.Mineralogical Society of America, pp. 351–403.

Borchert, H., Muir, R., 1964. Salt Deposits: The Origin, Metamorphism, andDeformation of Evaporites. D. Van Nostrand Co., Ltd., pp. 146–149.

Burns, R., 1993. Origin of electronic spectra of minerals in the visible to near-infraredregion. In: Pieters, C., Englert, P. (Eds.), Remote Geochemical Analysis: Elementaland Mineralogical Composition. Cambridge Univ. Press, New York, pp. 3–29.

Burns, R., Fisher, D., 1990. Iron–sulfur mineralogy of Mars: Magmatic evolution andchemical weathering products. J. Geophys. Res. 95, 14415–14421.

Chapman, M., Tanaka, K., 2001. Interior trough deposits on Mars: Subice volcanoes?J. Geophys. Res. 106 (E5), 10087–10100.

Chevrier, V., Mathe, P., Rochette, P., Grauby, O., Bourrie, G., Trolard, F., 2006. Ironweathering products in a CO2 (H2O or H2O2) atmosphere: Implications forweathering processes on the surface of Mars. Geochim. Cosmochim. Acta 70,4295–4317.

Christensen, P., and 15 colleagues, 2000. Detection of crystalline hematitemineralization on Mars by the Thermal Emission Spectrometer: Evidence fornear-surface water. J. Geophys. Res. 105, 9623–9642.

Christensen, P., and 25 colleagues, 2001a. Mars Global Surveyor Thermal EmissionSpectrometer experiment: Investigation description and surface science results.J. Geophys. Res. 106 (E10), 23823–23871.

Christensen, P., Morris, R., Lane, M., Bandfield, J., Malin, M., 2001b. Global mappingof martian hematite mineral deposits: Remnants of water-driven processes onearly Mars. J. Geophys. Res. 106 (E10), 23873–23885. doi:2000JE001415.

Christensen, P., and 26 colleagues, 2004. Mineralogy at Meridiani Planum from theMini-TES experiment on the Opportunity rover. Science 306, 1733–1739.

Clark, R., King, T., Klejwa, M., Swayze, G., Vergo, N., 1990. High resolution reflectancespectroscopy of minerals. J. Geophys. Res. 95 (B8), 12653–12680.

Clark, B., and 23 colleagues, 2005. Chemistry and mineralogy of outcrops atMeridiani Planum. Earth Planet. Sci. Lett. 240, 73–94.

Cloutis, E.and 11 colleagues, 2006. Detection and discrimination of sulfate mineralsusing reflectance spectroscopy. Icarus 157 (184), 121–157.

Gendrin, A., and 11 colleagues, 2005a. Identification of predominant ferricsignatures in association to the martian sulfate deposits. Lunar Planet. Sci.Conf. XXXVI. Abstract #1378.

Gendrin, A., and 10 colleagues, 2005b. Sulfates in martian layered terrains: TheOMEGA/Mars Express view. Science 307 (5715), 1587–1591. doi:10.1126/science.1109087.

Glotch, T., Christensen, P., 2005. Geologic and mineralogic mapping of Aram Chaos:Evidence for a water-rich history. J. Geophys. Res. 110, E09006. doi:10.1029/2004JE002389.

Glotch, T., Kraft, M., 2008. Thermal transformations of akaganéite and lepidocrociteto hematite: Assessment of possible precursors to martian crystalline hematite.Phys. Chem. Min. 35, 569–581.

Glotch, T., Rogers, A., 2007. Evidence for aqueous deposition of hematite- andsulfate-rich light-toned layered deposits in Aureum and Iani Chaos, Mars. J.Geophys. Res. 112, E06001. doi:10.1029/2006JE002863.

Glotch, T., Morris, R., Christensen, P., Sharp, T., 2004. Effect of precursor mineralogy onthe thermal infrared emission spectra of hematite: Application to martianhematite mineralization. J. Geophys. Res. 109, E07003. doi:10.1029/2003JE002224.

Glotch, T., Bandfield, J., Christensen, P., Calvin, M., McLennan, S., Clark, B., Rodgers, A.,Squyres, S., 2006. Mineralogy of the light-toned outcrop rock at Meridiani Planumas seen by the miniature Thermal Emission Spectrometer and implications for itsformation. J. Geophys. Res. 111, E12S03. doi:10.1026/2005HE002672.

Golden, D., Ming, D., Morris, R., Graff, T., 2008. Hydrothermal synthesis of hematitespherules and jarosite: Implications for diagnosis and hematite spherules insulfate outcrops at Meridiani Planum Mars. Am. Mineral. 93, 1201–1214.

Grotzinger, J., and 19 colleagues, 2005. Stratigraphy and sedimentology of a dry towet eolian depositional system, Burns Formation, Meridiani Planum, Mars.Earth Planet. Sci. Lett. 240, 11–72. doi:10.1016/j/epsl/2005.09.039.

Hoffman, N., 2001. Modern geothermal gradients on Mars and implications forsubsurface liquids. Conf. Geophys. Detection of Subsurface Water on Mars, p.7044.

Hunt, G., Salisbury, J., 1970. Visible and infrared spectra of minerals and rocks. I:Silicate minerals. Mod. Geol. 1, 283–300.

Hunt, G., Salisbury, J., Lenhoff, C., 1971. Visible and near infrared spectra of mineralsand rocks. III: Oxides and hydroxides. Mod. Geol. 2, 195–205.

Hutchison, L., Mustard, J., Gendrin, A., Bibring, J.-P., Langevin, Y., Gondet, B.,Mangold, N., and the OMEGA Science Team, 2005. Mafic polyhydrated sulfatesand kieserite in Capri Chasma. Lunar Planet. Sci. Conf. XXXVI. Abstract #1404.

Hynek, B., 2004. Extensive bedrock throughout Terra Meridiani, Mars: Implicationsfor hydrologic processes. Nature, 431. doi:10.1038/nature02902.

Hynek, B., Phillips, R., 2008. The stratigraphy of Meridiani Planum, Mars, andimplications for the layered deposits’ origin. Earth Planet. Sci. Lett. 274, 214–220.

Hynek, B., Arvidson, R., Phillips, R., 2002. Geologic setting and origin of TerraMeridiani hematite deposit on Mars. J. Geophys. Res., 107. doi:10.1029/2002JE001891.

ITT Visual Information Solutions, 2007. Using Continuum Removal. ENVI 4.4 Helpfile.

King, P., McSween, H., 2005. Effects of H2O, pH, and oxidation state on the stabilityof Fe minerals on Mars. J. Geophys. Res. 110, E12S10. doi:10.1029/2005JE002482.

Komatsu, G., Litasov, Y., 2002. Subice volcanism on the Azas Plateau: A comparisonwith possible Martian tuyas. Lunar Planet. Sci. Conf. XXXII. Abstract #1262.

Komatsu, G., Ori, G., Ciarcelluti, P., Litasov, Y., 2004. Interior Layered Deposit ofValles Marineris, Mars: Analogous subice volcanism related to Baikal Rifting,Southern Siberia. Planet. Space Sci. 52, 167–187. doi:10.1016/j.pss.2003.08.003.

Lane, M., Morris, R., Christensen, P., 1999. Spectral behavior of hematite at visible/near infrared and midinfrared wavelengths. In: The Fifth InternationalConference on Mars. California Institute of Technology, Pasadena, p. 6085.

Lane, M., Morris, R., Mertzman, S., Christensen, P., 2002. Evidence for platy hematitegrains in Sinus Meridiani, Mars. J. Geophys. Res. 107 (E12), 5126. doi:10.1029/2001JE001832.

Lane, M., Dyar, M., Bishop, J., 2004. Spectroscopic evidence for hydrous iron sulfatein the martian soil. Geophys. Res. Lett. 31, L19702. doi:10.1029/2004GL021231.

Le Deit, L., Le Mouélic, S., Bourgeois, O., Combe, J.-P., Mège, D., Sotin, C., Gendrin, A.,Hauber, E., Mangold, N., Bibring, J.-P., 2008. Ferric oxides in East CandorChasma, Valles Marineris (Mars) inferred from analysis of OMEGA/Mars Expressdata: Identification and geological interpretation. J. Geophys. Res. 113, E07001.doi:10.1029/2007JE002950.

Lichtenberg, K., Arvidson, R., Morris, R., Murchie, S., Bishop, J., Glotch, T., Dobrea, E.,Mustard, J., Andrews-Hanna, J., Roach, L., 2009. Stratigraphy and relationship ofhydrated minerals in the layered deposits of Aram Chaos, Mars. Lunar Planet.Sci. Conf. XL. Abstract #2326.

Lucchitta, B., 1981. Valles Marineris – Faults, volcanic rocks, channels, basin beds.Reports of Planetary Geology Program, NASA-TM 84211, pp. 419–421.

Lucchitta, B., 1982. Lakes or playas in Valles Marineris. Reports of Planetary GeologyProgram, NASA-TM 85127, pp. 233–234.

Lucchitta, B., McEwen, A., Clow, G., Geissler, P., Singer, R., Schultz, R., Squyres, S.,1992. The Canyon system on Mars. In: Kieffer, H., Jakosky, B., Snyder, C.,Matthews, M. (Eds.), Mars. Univ. Arizona Press, Tucson, pp. 453–492.

Lucchitta, B., Isbell, N., Howington, A., 1994. Topography of Valles Marineris:Implications for erosional and structural history. J. Geophys. Res. 99 (E2), 3783–3798.

Malin, M., and 13 colleagues, 2007. Context camera investigation on board the MarsReconnaissance Orbiter. J. Geophys. Res. 112, E05S04. doi:10.1029/2006JE002808.

Mangold, N., Gendrin, A., Gondet, B., Lemouelic, S., Quantin, C., Ansan, V., Bibring,J.-P., Langevin, Y., Masson, P., Neukum, G., 2008. Spectral and geological study ofthe sulfate-rich region of West Candor Chasma, Mars. Icarus 194, 519–543.

McCauley, J., 1978. Atlas of Mars, USGS Misc. Inv. Serv. Map I-897 (MC-18).McCollom, T.M., Hynek, B.M., 2005. A volcanic environment for bedrock diagenesis

at Meridiani Planum on Mars. Nature 438, 1129–1131. doi:10.1038/nature04390.

McCord, T., Huguenin, R., Mink, D., Pieters, C., 1977. Spectral reflectance of martianareas during the 1973 opposition: Photoelectric filter photometry 0.33–1.10 lm. Icarus 31, 25–39.

McEwen, A., and 14 colleagues, 2005. Mars Reconnaissance Orbiter’s HighResolution Imaging Science Experiment (HiRISE). J. Geophys. Res. 11, E05S02.doi:10.1029/2005JE002605.

McLennan, S.M., and 31 colleagues, 2005. Provenance and diagenesis of theevaporite-bearing Burns formation, Meridiani Planum, Mars. Earth Planet. Sci.Lett. 240 (1), 95–121. doi:10.1016/j.epsl.2005.09.041.

Morris, R., Lauer, H., Lawson, C., Gibson, E., Nace, G., Stewart, C., 1985. Spectral andother physicochemical properties of submicron powders of hematite (a-Fe2O3),maghemite (c-Fe2O3), magnetite (Fe3O4), goethite (a-FeOOH), and lepidocrocite(c-FeOOH). J. Geophys. Res. 90, 3126–3144.

Morris, R.V., Golden, D.C., Bell III, J.F., 1997. Low-temperature reflectivity spectra ofred hematite and the color of Mars. J. Geophys. Res. 102, 9125–9133.doi:10.1029/96JE03993.

Morris, R.V., and 11 colleagues, 2000. Mineralogy, composition, and alteration ofMars Pathfinder rocks and soils: Evidence from multispectral, elemental, andmagnetic data on terrestrial analogue, SNC meteorite, and Pathfinder samples.J. Geophys. Res. 105, 1757–1817. doi:10.1029/1999JE001059.

Morris, R.V., and 23 colleagues, 2006. Mossbauer mineralogy of rock, soil, and dustat Meridiani Planum, Mars: Opportunity’s journey across sulfate-rich outcrop,basaltic sand and dust, and hematite lag deposits. J. Geophys. Res. 112, E05S11.doi:10.1029/2006JE002682.

Murchie, S.M., and 49 colleagues, 2007. Compact Reconnaissance ImagingSpectrometer for Mars (CRISM) on Mars Reconnaissance Orbiter (MRO). J.Geophys. Res. 112, E05S03. doi:10.1029/2006JE002682.

Murchie, S., and 12 colleagues, 2009a. Evidence for the origin of layered deposits inCandor Chasma, Mars, from mineral composition and hydrologic modeling. J.Geophys. Res. 114, E00D05. doi:10.1029/2009JE003343.

Murchie, S., and 19 colleagues, 2009b. The CRISM investigation and data set fromthe Mars Reconnaissance Orbiter’s primary science phase. J. Geophys. Res. 114,E00D07. doi:10.1029/2009JE003344.

Mustard, J., Bell, J., 1994. New composite reflectance spectra of Mars from 0.4 to3.14 lm. Geophys. Res. Lett. 21, 353–356.

674 L.H. Roach et al. / Icarus 207 (2010) 659–674

Mustard, J., Poulet, F., Gendrin, A., Bibring, J.-P., Langevin, Y., Gondet, B., Mangold, N.,Bellucci, G., Altieri, F., 2005. Olivine and pyroxene diversity in the crust of Mars.Science 307, 1594–1597.

Nedell, S., Squyres, S., Anderson, D., 1987. Origin and evolution of the layereddeposits in the Valles Marineris, Mars. Icarus 70, 409–441.

NoeDobrea, E., Poulet, F., Malin, M., 2008. Correlations between hematite andsulfates in chaotic terrain east of Valles Marineris. Icarus 193, 516–534.

Peterson, C., 1981. A secondary origin for the central plateau of Hebes Chasma. Proc.Lunar Planet. Sci. Conf. 12, 1459–1471.

Pinet, P., Chevrel, S., 1990. Spectral identification of geological units on the surfaceof Mars related to the presence of silicates from Earth-based near-infraredtelescopic charge-coupled device imaging. J. Geophys. Res. 95, 14435–14446.

Poulet, F., Gomez, C., Bibring, J.-P., Langevin, Y., Gondet, B., Pinet, P., Belluci, G.,Mustard, J., 2007. Martian surface mineralogy from Observatoire pour laMineralogie, l’Eau, les Glaces et l’Activite on board the Mars Express spacecraft(OMEGA/MEx): Global mineral maps. J. Geophys. Res., 112. doi:10.1029/2006JE002840.

Poulet, F., Arvidson, R., Gomez, C., Morris, R., Bibring, J.-P., Langevin, Y., Gondet, B.,Griffes, J., 2008. Mineralogy of Terra Meridiani and western Arabia Terra fromOMEGA/MEx and implications for their formation. Icarus 195, 106–130.

Putzig, N., Mellon, M., 2007. Apparent thermal inertia and the surface heterogeneityof Mars. Icarus 191, 68–94.

Roach, L.H., Mustard, J.F., Murchie, S.L., Bishop, J.L., Weitz, C.M., Knudson, A.T.,Bibring, J.-P., Pelkey, S.M., Ehlmann, B.L., and the CRISM Science Team, 2007.Magnesium and iron sulfate variety and distribution in East Candor and CapriChasma, Valles Marineris. In: 7th Int. Conf. on Mars, p. 2332.

Roach, L.H., Mustard, J.F., Murchie, S.L., Bibring, J.-P., Forget, F., Lewis, K.W., Aharonson,O., Vincendon, M., Bishop, J.L., 2009a. Testing evidence of recent hydration statechange in sulfates on Mars. J. Geophys. Res.. doi:10.1029/2008JE003245.

Roach, L.H., Mustard, J.F., Swayze, G., Milliken, R.E., Bishop, J.L., Ehlmann, B.L.,Murchie, S.L., Lichtenberg, K., 2009b. Hydrated mineral stratigraphy of IusChasma, Valles Marineris. Icarus 206, 253–268.

Schultz, R., 1998. Multiple-process origin of Valles Marineris basins and troughs,Mars. Planet. Space Sci. 46, 827–829. doi:10.1016/S0032-0633(98)00030-0.

Schultz, R.A., Lin, J., 2001. Three-dimensional normal faulting models of the VallesMarineris, Mars, and geodynamical implications. J. Geophys. Res. 106, 16549–16566.

Sefton-Nash, E., Catling, D.C., 2008. Hematitic concretions at Meridiani Planum,Mars: Their growth timescale and possible relationship with iron sulfates. EarthPlanet. Sci. Lett. 269, 366–376. doi:10.1016/j.epsl.2008.02.009.

Sonnenfeld, P., 1984. Brines and Evaporites. Academic Press, New York. pp. 291–297.

Squyres, S.W., Knoll, A.H. (Eds.), 2005. Sedimentary Geology at Meridiani Planum,Mars. Elsevier, Amsterdam.

Squyres, S.W., and 53 colleagues, 2006. Overview of the opportunity Marsexploration rover mission to Meridiani Planum: Eagle Crater to PurgatoryRipple. J. Geophys. Res. 111, E12S12. doi:10.1029/2006JE002771.

Stewart, F.H., 1956. Replacements involving early carnallite in the potassium-bearing evaporites of Yorkshire. Mineral. Mag. 31 (233), 127–135. doi:10.1180/minmag.1956.031.233.02.

Tosca, N.J., McLennan, S.M., Clark, B.C., Grotzinger, J.P., Hurowitz, J.A., Knoll, A.H.,Schroder, C., Squyres, S.W., 2005. Geochemical modeling of evaporationprocesses on Mars: Insight from the sedimentary record at Meridiani Planum.Earth Planet. Sci. Lett. 240 (1), 122–148. doi:10.1016/j.epsl.2005.09.042.

Tosca, N.J., Milliken, R.E., Wyatt, M.B., 2008. An experimental approach toclay mineral formation on Mars. Geochim. Cosmochim. Acta 72 (Suppl. 1),952.

Vaniman, D.T., Chipera, S.J., 2006. Transformations of Mg- and Ca-sulfate hydratesin Mars regolith. Am. Mineral. 91, 1628–1642.

Weitz, C.M., 1999. A volcanic origin for the Interior Layered Deposits in HebesChasma, Mars. Lunar Planet. Sci. XXX. Abstract #1736.

Weitz, C., Lane, M., Staid, M., Noe Dobrea, E., 2008. Gray hematite distribution andformation in Ophir and Candor Chasmata. J. Geophys. Res. 113. doi:10.1029/2007je002930.

Witbeck, N., Tanaka, K., Scott, D., 1991. Geologic map of the Valles Marineris region,Mars, scale 1:2000000. US Geol. Surv. Misc. Inv. Series. 1-2010.

Wolff, M., Lee, S., Clancy, R., Margin, L., Bell, J., James, P., 1997. 1995 observations ofMartian dust activity using the Hubble Space Telescope. J. Geophys. Res. 104,8987–9007.

Yen, A., and 35 colleagues, 2005. An integrated view of the chemistry andmineralogy of martian soils. Nature 436, 49–54. doi:10.1038/nature03637.

Zuber, M., Smith, D., Solomon, S., Muhleman, D., Head, J., Garvin, J., Abshire, J.,Bufton, J., 1992. The Mars Observer laser altimeter investigation. J. Geophys.Res. 97 (5), 7781–7797.