Early Mars Environments Revealed Through Near-Infrared Spectroscopy of Alteration Minerals

Home , Arnold (crater), Baldet (Martian crater), Bethany Ehlmann, Billy (crater), NASA Academy, Sinton (crater)

Bethany List Ehlmann A.B., Washington University in St. Louis, 2004 M.Sc., University of Oxford, 2005 MSc. (by research), University of Oxford, 2007 Sc.M., Brown University, 2008

A dissertation

Submitted in partial fulfillment of the requirements for the degree of Doctor of Philosophy in the Program in Geological Sciences at Brown University

Providence, Rhode Island

May 2010

Curriculum Vitae

Bethany L. Ehlmann Dept. of Geological Sciences [email protected] Box 1846, Brown University (office) +1 401.863.3485 (mobile) +1 401.263.1690 Providence, RI 02906 USA

EDUCATION Ph.D. Brown University, Planetary Sciences, anticipated May 2010 Advisor: Prof. John Mustard Sc. M., Brown University, Geological Sciences, 2008 Advisor: Prof. John Mustard M.Sc. by research, University of Oxford, Geography (Geomorphology), 2007 Advisor: Prof. Heather Viles M.Sc. with distinction, University of Oxford, Environmental Change & Management, 2005 Advisor: Dr. John Boardman A.B. summa cum laude, Washington University in St. Louis, 2004 Majors: Earth & Planetary Sciences, Environmental Studies; Minor: Mathematics Advisor: Prof. Raymond Arvidson International Baccalaureate Diploma, Rickards H.S., Tallahassee, Florida, 2000

Additional Training: Nordic/NASA Summer School: Water, Ice and the Origin of Life in the Universe, Iceland, July 2009 Vatican Observatory Summer School in Astronomy &Astrophysics, Castel Gandolfo, Italy, 2005 Rainforest to Reef Program: Marine Geology, James Cook University, Australia, July 2004 School for International Training, Development and Conservation Program, Panamá, Sept-Dec 2002

PROFESSIONAL EXPERIENCE Planetary Scientist III, Jet Propulsion Laboratory, accepted position, start 9/2011 Marie Curie Incoming International Fellow, Institut d’Astrophysique Spatiale, France, start 7/2010 Compact Reconnaissance Imaging Spectrometer for Mars (CRISM) Science Team Collaborator, 9/2006-present Graduate Research Fellow, Brown University, 9/2006-present Postgraduate Researcher, School of Geography & Environment, Environmental Change Institute, University of Oxford, 2004-2006. Mars Exploration Rovers (MER) Athena Science Team, Science Collaborator, 5/2003-9/2004 Undergraduate Researcher, Remote Sensing Laboratory, Washington University, 2001-2004 Space Studies Board Intern, National Research Council, National Academy of Sciences, 2003 Student Science Consultant, Interdisciplinary Enviro. Law Clinic, Washington Univ. School of Law, 2003 Research Associate, NASA Astrobiology Academy, Ames Research Center, 2002

AWARDS AND FELLOWSHIPS Sherwood Chang-Eliot Kalmbach Award for Excellence in Astrobiology Research, student speaker award at the Gordon Origins of Life conference (2010) Best student oral presentation, runner-up, International Clay Conference, Italy (2009) Pellas-Ryder Award for best peer-reviewed planetary sciences paper authored by a student, Geological Society of America and Meteoritical Society (2009) Best Student Paper Award, Planetary Sciences Section, American Geophys. Union Fall Meeting (2008) NASA Group Achievement Award, Mars Exploration Rover Science Operations Team (2005) National Science Foundation Graduate Research Fellowship (2004-2009) Rhodes Scholar (Missouri and Keble/Hertford, 2004) Morris K. Udall Scholar in Environmental Studies (2002, 2003) Barry M. Goldwater Scholar in Science, Mathematics, and Engineering (2002)

PEER-REVIEWED PUBLICATIONS 31. Ehlmann, BL, JF Mustard, RN Clark, GA Swayze, SL Murchie. Evidence for low-grade metamorphism, diagenesis, and hydrothermal alteration on Mars from phyllosilicate mineral assemblages. Clays & Clay Minerals, in review. 30. Viles, HA, BL Ehlmann, CF Wilson, T Cebula, M Bourke. Simulating physical weathering of basalt on Mars and Earth by thermal cycling, Geophys. Res. Lett., in review 29. Wray, JJ, Ehlmann, BL. Morphology and mineralogy of possible Martian methane source regions, Planetary & Space Sci., in review 28. Ehlmann, BL, JF Mustard, S.L. Murchie. Geologic setting of serpentine-bearing rocks on Mars. Geophys. Res. Letters, doi:10.1029/2010GL042596, 2010. 27. Skok, JR, JF Mustard, SL Murchie, MB Wyatt, BL Ehlmann. Spectrally distinct ejecta in Syrtis Major, Mars: Evidence for environmental change at the Hesperian-Amazonian boundary. J. Geophys. Res., 115, E00D14, doi:10.1029/2009JE003338, 2010. 26. Ehlmann, BL. Diverse aqueous environments during Mars' first billion years: the emerging view from orbital visible-near infrared spectroscopy. Geochemical News, 142, 2010. 25. Mustard, JF, BL Ehlmann, F Poulet, N Mangold, JW Head, SL Murchie, J-P Bibring, LH Roach. Composition, Morphology, and Stratigraphy of Noachian/Phyllosian Crust around the Isidis basin. J. Geophys. Res, 114, E00D12, doi:10.1029/2009JE003349. 24. McKeown, N. K., J. L. Bishop, E. Z. Noe Dobrea, B. L. Ehlmann, M. Parente, J. F. Mustard, S. L. Murchie, G. A. Swayze, J. Bibring, and E. A. Silver. Characterization of phyllosilicates observed in the central Mawrth Vallis region, Mars, their potential formational processes, and implications for past climate, J. Geophys. Res., 114, E00D10, doi:10.1029/2008JE003301, 2009. 23. Ehlmann, BL, JF Mustard, GA Swayze, RN Clark, JL Bishop, F Poulet, D Des Marais, LH Roach, RE Milliken, J Wray, O Barnouin-Jha SL Murchie. Identification of hydrated silicate minerals on Mars using MRO-CRISM: geologic context near Nili Fossae and implications for aqueous alteration, J. Geophys. Res, E00D08, doi:10.1029/2009JE003339, 2009. 22. Murchie, SL, JF Mustard, BL Ehlmann, RE Milliken, JL Bishop, NK McKeown, EZ Noe Dobrea, FP Seelos, DL Buczkowski, SM Wiseman, RE Arvidson, JJ Wray, G Swayze, RN Clark, J-P Bibring, AS McEwen. A synthesis of Martian aqueous mineralogy after one Mars year of observations from the Mars Reconnaissance Orbiter. J. Geophys. Res., doi:10.1029/2009JE003344, 2009. 21. Ehlmann, BL, JF Mustard, SL Murchie, F Poulet, JL Bishop, AJ Brown, WM Calvin, RN Clark, DJ Des Marais, RE Milliken, LH Roach, TL Roush, GA Swayze, JJ Wray. Orbital Identification of Carbonate-Bearing Rocks on Mars. Science, 322, 1828-1832, 2008 20. Milliken, RE, G Swayze, R Arvidson, J Bishop, R Clark, B Ehlmann, R Green, J Grotzinger, R Morris, S Murchie, J Mustard, C Weitz. Opaline silica in (geologically?) young deposits on Mars. Geology, 36(11), 847-850, 2008. 19. Bishop, JL, EZ Noe Dobrea, NK McKeown, M Parente, BL Ehlmann, JR Michalski, RE Milliken, F Poulet, GA Swayze, JF Mustard, SL Murchie, J-P Bibring. Phyllosilicate Diversity and Past Aqueous Activity Revealed at Mawrth Vallis, Mars, Science 321, 830-833, 2008. 18. Wray, J.J., BL Ehlmann, SW Squyres, JF Mustard, RL Kirk. Compositional Stratigraphy of Clay- Bearing Layered Deposits at Mawrth Vallis, Mars. Geophysical Research Letters 35, L12202, doi: 10.1029/2008GL034385, 2008. 17. Mustard JF, SL Murchie, SL, SM Pelkey, BL Ehlmann, RE Milliken, JA Grant, J-P Bibring, F Poulet, J Bishop, E Noe Dobrea, L Roach, F Seelos, RE Arvidson, S Wiseman, R Green, C Hash, D Humm, E Malaret, JA McGovern, K Seelos, T Clancy, R Clark, D Des Marais, N Izenberg, A Knudson, Y Langevin, T Martin, P McGuire, R Morris, M Robinson, T Roush, M Smith, G Swayze, H Taylor, T Titus, M Wolff. Hydrated Silicate Minerals on Mars Observed by the CRISM Instrument on MRO. Nature 454, 305-309, 2008. 16. Herkenhoff, KE and 44 others including (B Ehlmann) Surface processes recorded by rocks and soils on Meridiani Planum, Mars: Microscopic Imager observations during Opportunity's first three extended missions, J. Geophys. Res., 113, E12S32, doi:10.1029/2008JE003100. 15. Ehlmann, BL, JF Mustard, CI Fassett, SC Schon, JW Head, DJ Des Marais, JA Grant, SL Murchie, CRISM team. Clay mineralogy and organic preservation potential of lacustrine sediments from a Martian delta environment, Jezero Crater, Nili Fossae, Mars. 2008. Nature Geoscience, 1, 355-358, 2008. 14. Ehlmann, BL, HA Viles, and MC Bourke. Quantitative morphologic analysis of boulder shape and surface texture to infer environmental history: A case study of rock breakdown at the Ephrata Fan, Channeled Scabland, Washington, J. Geophys. Res., 113, F02012, doi:10.1029/2007JF000872, 2008.

13. Herkenhoff, KE and 42 others (including B.L. Ehlmann) Overview of the Microscopic Imager Investigation during Spirit’s first 450 sols in Gusev crater. J. Geophys. Res., 111, E02S04, doi:10.1029/2005JE002574. 12. Ehlmann, BL and RE Criss. Enhanced Stage and Stage Variability on the Lower Missouri River benchmarked by Lewis and Clark, Geology 34(11), 977-980, 2006. 11. Arvidson, RE, F Poulet, RV Morris, J-P Bibring, JF Bell III, SW Squyres, PR Chistensen, G Belluci, B Gondet, BL Ehlmann, WH Farrand, RL Fergason, M Golombek, JL Griffes, J Grotzinger, EA Guinness, KE Herkenhoff, JR Johnson, G Klingelhoefer, Y Langevin, D Ming, K Seelos, RJ Sullivan, JG Ward, SM Wiseman, M Wolff. Nature and Origin of the Hematite-Bearing Plains of Meridiani Based on Analyses of Orbital and Opportunity Data Sets, JGR 111(E12), E12S08, doi:10.1029/2006JE002728, 2006. 10. Herkenhoff, KE, SW Squyres, R Anderson, BA Archinal, RE Arvidson, JM Barrett, KJ Becker, JF Bell III, C Budney, NA Cabrol, MG Chapman, D Cook, BL Ehlmann, J Farmer, B Franklin, LR Gaddis, DM Galuszka, PA Garcia, TM Hare, E Howington-Kraus, JR Johnson, S Johnson, K Kinch, RL Kirk, EM Lee, C Leff, M Lemmon, MB Madsen, JN Maki, KF Mullins, BL Redding, L Richter, MR Rosiek, MH Sims, LA Soderblom, N Spanovich, R Springer, RM Sucharski, T Sucharski, R Sullivan, JM Torson, A Yen. Overview of the Microscopic Imager Investigation during Spirit’s first 450 sols in Gusev crater. J. Geophys. Res. 111, doi: 10.1029/2005JE002574, 2006. 9. Ehlmann, BL, RE Arvidson, BL Jolliff, SS Johnson, B Ebel, N Lovenduski, JD Morris, JA Byres, NO Snider, RE Criss. Hydrologic and Isotopic Modeling of Alpine Lake Waiau, Mauna Kea, Hawaii. Pacific Science 59 (1), 1-15, 2005. 8. Ehlmann, BL, J Chowdhury, TC Marzullo, RE Collins, J Litzenberger, S Ibsen, WR Krauser, B DeKock, M Hannon, J Kinnevan, R Shepard, FD Grant. Humans to Mars: A Feasibility and Cost- Benefit Analysis. Acta Astronautica 56 (9-12), 851-858, 2005. 7. Golombek, MP, RE Arvidson, JF Bell III, PR Christensen, JA Crisp, LS Crumpler, BL Ehlmann, RL Fergason, JA Grant, R Greeley, AFC. Haldemann, DM Kass, TJ Parker, JT Schofield, SW Squyres, RW Zurek. Assessment of Mars Exploration Rover Landing Site Predictions and Implications for Climate Change. Nature, 436, 43-46, 2005. 6. Arvidson, RE, RC Anderson, P Bartlett, JF Bell III, PR Christensen, P Chu, K Davis, BL Ehlmann, MP Golombek, S Gorevan, EA Guinness, AFC. Haldemann, KE Herkenhoff, G Landis, R Li, R Lindemann, DW Ming, T Myrick, T Parker, L Richter, FP Seelos IV, LA Soderblom, SW Squyres, RJ Sullivan, J Wilson. Localization and Physical Properties Experiments Conducted by Opportunity at Meridiani Planum. Science 306, 1730-1733, 2004. 5. Arvidson, RE , RC Anderson, P Bartlett, JF Bell III, D Blaney, PR Christensen, P Chu, L Crumpler, K Davis, BL Ehlmann, R Fergason, MP Golombek, S Gorevan, JA. Grant, R Greeley, EA Guinness, AFC. Haldemann, KE Herkenhoff, J Johnson, G Landis, R Li, R Lindemann, H McSween, DW Ming, T Myrick, L Richter, FP Seelos IV, SW Squyres, R Sullivan, A Wang, J Wilson. Localization and physical properties experiments conducted by Spirit at Gusev crater. Science 305, 821-824, 2004. 4. Grant, JA , R Arvidson, JF Bell, III, NA Cabrol, MH Carr, P Christensen, L Crumpler, DJ Des Marais, BL Ehlmann, J Farmer, M Golombek, FD Grant, R Greeley, K Herkenhoff, R Li, HY McSween, DW Ming, J Moersch, JW Rice, Jr., S Ruff, L Richter, S Squyres, R Sullivan, C Weitz. Surficial Deposits at Gusev Crater Along Spirit Rover Traverses. Science 305, 807-810, 2004. 3. Herkenhoff, KE, SW Squyres, R Arvidson, DS Bass, JF Bell III, P Bertelsen, BL Ehlmann, W Farrand, L Gaddis, R Greeley, J Grotzinger, AG Hayes, SF Hviid, JR Johnson, B Jolliff, KM Kinch, AH Knoll, MB Madsen, JN Maki, SM McLennan, HY McSween, JW Rice, Jr., L Richter, M Sims, PH Smith, LA Soderblom, N Spanovich, R Sullivan, S Thompson, T Wdowiak, C Weitz, P Whelley. Evidence from Opportunity’s Microscopic Imager for Ancient Water on Meridiani Planum. Science 306, 1727-1730, 2004. 2. Soderblom, LA, RC Anderson, RE Arvidson, JF Bell III, NA Cabrol, W Calvin, PR Christensen, BC Clark, T Economou, BL Ehlmann, WH Farrand, D Fike, R Gellert, TD Glotch, MP Golombek, R Greeley, JP Grotzinger, KE Herkenhoff , DJ Jerolmack, JR Johnson, B Jolliff , G Klingelhöfer, AH Knoll, ZA Learner, R Li, MC Malin, SM McLennan, HY McSween, DW Ming, RV Morris, JW Rice Jr., L Richter, R Rieder, D Rodionov, C Schröder, FP Seelos IV, JM Soderblom, SW Squyres, R Sullivan, WA Watters, CM Weitz, MB Wyatt, A Yen, J Zipfel. Soils of Eagle Crater and Meridiani Planum at the Opportunity Rover Landing Site. Science 306, 1723-1726, 2004. 1. Solar System Exploration Survey, National Academy of Sciences, Space Studies Board. (Brian Dewhurst, Bethany Ehlmann, and David Smith, text ed.), New Frontiers in Solar System Exploration. National Academies Press, 2003

FIRST-AUTHOR CONFERENCE PUBLICATIONS AND PRESENTATIONS Ehlmann, BL, JF Mustard. “Stratigraphy of the Nili Fossae and the Jezero crater watershed: a reference section for the Martian clay cycle.” First International Conference on Mars Sedimentology and Stratigraphy, El Paso, Texas, April 19-21, 2010, Abstract #6064 (talk). Ehlmann, BL, JF Mustard, DL Bish. “Weathering and hydrothermal alteration of basalts in Iceland: mineralogy from VNIR, TIR, XRD and implications for linking Mars orbital and surface datasets.” LPSC 41, Houston, Texas, March 1-5, 2010, Abstract #1858 (poster). Ehlmann, BL, JF Mustard, SL Murchie. “Geologic setting of serpentine deposits on Mars.” LPSC 41, Houston, Texas, March 1-5, 2010, Abstract #2235 (talk). (invited) Ehlmann, BL, JF Mustard. “Evidence for diverse hydrothermal and weathering environments on ancient (>3.5Ga) Mars from recent orbital observations” Gordon Research Conference on The Origin of Life, Galveston, TX, January 2010 (talk) Ehlmann, BL, JF Mustard, DL Bish. “Weathering and hydrothermal alteration of basalts in Iceland: mineralogy from VNIR, TIR, XRD and implications for Mars.” Am. Geophys. Union Fall Mtg. 2009. (poster) Ehlmann, BL, JF Mustard, SL Murchie. “Extensive aqueous alteration of Mars’ earliest crust: recent results from NASA’s CRISM hyperspectral imager and implications for planetary habitability.” Vernadsky/Brown Microsymposium on Comparative Planetology. Moscow, Russia, October 2009. (talk) Ehlmann, BL and JF Mustard. “Regional carbonate- and kaolinite-bearing rock units and how precusros lithologies control alteration products on Mars: An example from the Nili Fossae Region.” New Martian Chemistry Workshop, Medford, Massachusetts, July 27-28, 2009, Abstract #8018 (talk) Ehlmann, BL et al. “Evidence for Low-Grade Metamorphism/Diagenesis on Mars from Phyllosilicate Mineral Assemblages” 14th International Clay Conference, Castellaneta Marina, Italy, June 14-20, 2009, Abstract #MC1c.L3 (talk) Ehlmann, BL, et al. “Detection of Serpentine on Mars by MRO-CRISM and Possible Relationship with Olivine and Magnesium Carbonate in Nili Fossae” LPSC 40, Houston, Texas, March 21-25, 2009, Abstract #1787 (talk) Ehlmann, BL, et al. “Modeling Modal Mineralogy of Laboratory Mixtures of Nontronite and Mafic Minerals form Visible Near-Infrared Spectral Data” LPSC 40, Houston, Texas, March 21-25, 2009, Abstract #1771 (poster) (invited) Ehlmann, BL et al. “Orbital Identification of Carbonate-Bearing Rocks on Mars” Am. Geophys. Union Fall Mtg. 2008 (talk) (invited) Ehlmann, BL et al. “Diverse Alteration Minerals Around Martian Impact Craters Revealed by MRO-CRISM: Indicators of Hydrothermal Activity or Subsurface Aqueous Alteration?” Am. Geophys. Union Fall Mtg. 2008 (talk) Ehlmann, BL et al. “Phyllosilicates, zeolites, and carbonate near Nili Fossae: evidence for distinct environments of aqueous alteration.” Workshop on Martian Phyllosilicates, Paris, October 2008. (talk) Ehlmann, BL, et al. “Distinct provinces of aqueous alteration in the western Isidis region identified with MRO-CRISM” LPSC 39, Houston, Texas, March 10-14, 2008, Abstract #2326 (talk) Ehlmann, BL, et al. “Infrared spectra of impact products from Lonar Crater: the effects of weathering and implications for Mars.” LPSC 39, Houston, Texas, March 10-14, 2008, Abstract #2437 (poster) Ehlmann, BL, et al. “Mineralogic diversity and geomorphology of CRISM-detected phyllosilicate bearing materials in Nili Fossae, Mars: Implications for aqueous alteration.” Eos Trans. Am. Geophys. Union 88(52), 2007 Fall Meet. Suppl., Abstract H43H-01 (talk). Ehlmann, BL, et al. “New secondary minerals detected by MRO CRISM and their geologic settings: kaolinite, chlorite, illite/muscovite, and the possibility of serpentine or carbonate in Nili Fossae” Seventh International Conference on Mars, Pasadena, California, July 9-13, 2007, Abstract #3270. (talk) Ehlmann, BL, et al. “New phyllosilicate mineral signatures from west of Nili Fossae, Mars through combined OMEGA-CRISM analysis” LPSC 38, Houston, Texas, March 12-15, 2007, Abstract #2078. (poster) Ehlmann, BL, et al. “Quantifying boulder shape and surface texture: A case study using geomorphology to infer environmental history at the Ephrata Fan, Channeled Scablands, Washington” LPSC 38, Houston, Texas, March 12-15, 2007, Abstract #1325. (poster) Ehlmann, BL and RE Criss. “Enhanced Stage Variability on the Lower Missouri River as benchmarked by Lewis and Clark: Implications for Ecosystem Restoration” Eos Trans. Am. Geophys. Union 87(52), 2006 Fall Meet. Suppl., Abstract H43H-01 (talk).

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Ehlmann, BL and HA Viles. “Fluvial feature persistence and lichen weathering rates on basalt boulders, Ephrata Fan, Washington” Geomorphology & Earth Systems Science, BGRG International Conference, Loughborough, UK, June 28-30, 2006 (poster). Ehlmann, BL et al., “Terrain Roughness from MER Traverse Profiles at Gusev Crater and Meridiani Planum” Eos Trans. AGU, 85(47), 2004 Fall Meet. Suppl., Abstract P21A-0201. (poster) Sullivan, R. et al., 2004. (talk by B. Ehlmann) “Rock and Soil Physical Properties at the MER Terra Meridiani Landing Site”. European Geophysical Union Meeting, Nice, France, April 25-30, 2004 Ehlmann, BL and RE Criss. “Stage Variability of the Missouri River as recorded by Lewis and Clark.” Geological Society of America Abstracts with Programs, Vol. 36, No. 3, p. 7, GSA Regional Meeting, St. Louis, April 1-2, 2004. (talk) Ehlmann, BL et al., 2002. “Hydrology of Lake Waiau” Eos. Trans. AGU, 83(19), 2002 Spring Meet. Suppl., Abstract H51E-06. (talk)

TEACHING EXPERIENCE Teaching Assistant, GE133, Remote Sensing of the Environment, Brown University, 2010. Sherdian Center Certificate IV: Teaching Consultant Program, teaching consultant 2009-2010. Teaching assistant, geosciences undergraduate spring break field trip, Brown University 2008 Sherdian Center Certificate I: Introduction to Reflective Teaching Practice, five course seminar, May 2007 Tutor for Remote Sensing-GIS, Oxford University Hilary and Trinity terms 2005, Michaelmas term 2006. Hired by colleges to teach seminar courses in remote sensing (term-long and revision) for first-year Geography students (Mansfield, Worchester, St. Hilda’s, Christ Church, Merton, Wadham, St. Edmond’s Hall Colleges) Teaching Assistant, EPSc 407 Remote Sensing, Washington University, 2003 Honorable Mention, Best Teaching Assistant, Earth & Planetary Sciences Dept., May 2003 Academic Tutor, Earth & Planetary Sci. Courses, Student Educational Services, Washington Univ., 2001-2002

TALKS AND PUBLIC OUTREACH 4/10 "Aqueous environments during Mars' first billion years ", invited talk for the D. Foster Hewitt Lecture Series, Lehigh University 11/09 "Aqueous environments during Mars' first billion years ", invited talk at Washington University, St. Louis 11/09 “Following the water and hunting for habitats on Mars: new results from the last 5 years of NASA exploration missions” Indiana University, undergraduate geosciences group 09/09 "Aqueous environments during Mars' first billion years ", invited talks at Caltech, Jet Propulsion Laboratory 03/09 “Evidence for Hydrothermal Environments on Early Mars”, talk at Space Sciences Branch, NASA-Ames 01/09 Curriculum preparation for the National Science Teachers Association. “What we can learn at different spectral and spatial resolutions: an example from Mars” 12/08 Panelist in NASA press conference at AGU; interviews with Canadian Television (CTV), BBC radio, Earth & Sky, SETI Institute Science radio; prepared press release images for Mars carbonate discovery 08/08 Assisted with Brown University blogs from the field in Iceland by Richard Lewis, Brown Media Services 07/08 Blog interview with The Scientific Activist, new findings about Water on Mars (http://scienceblogs.com/scientificactivist/2008/07/interview_with_a_planetary_geo.php) 06/08 Interviewed for discovery of clays in ancient Mars lake deposits (The Guardian, National Geographic, Associated Press, San Francisco Chronicle); prepared press release images 9/07-4/08 Mentor for JHU-APL’s CRISM Mars Exploration Student Data Team, Kickapoo H.S., Missouri 2/08 “Composition and stratigraphy of altered Noachian crust: new results from MRO CRISM for phyllosilicates on Mars”, invited talk at the Planetary Sciences Institute 3/08 Planetary Society guest blogger for the Lunar and Planetary Sciences Conference 2008 8/07 Docent for RI Museum of Natural History “Mars 3-D” exhibit 7/07 NASA Ames Academy for Exploration talk “New Insights into Water on Mars: First results from CRISM” viii

5/07 Rickards High School International Baccalaureate Class Senior Celebration keynote speaker “The importance of breadth and internationalism in education” 9/06-01/10 bi-semesterly Earth science lessons for 2nd & 4th graders at Vartan Gregorian Elementary School 2/05 Oxford University Space and Astronomical Society talk “First Year on Mars: Science Results from the NASA Mars Exploration Rovers” 9/05 Edwardsville Middle School “NASA Mars Exploration Rover Mission” 8/04 Macquarie University, Australian Centre for Astrobiology invited talk “First Results from the Mars Exploration Rovers at Gusev Crater and Meridiani Planum” 3/04 Panelist in two live, televised weekly JPL Mars Exploration Rover press conferences 1/04 Guest on Fox 2 News in the Morning, live (KTVI-St. Louis) : Mars Rovers Landings

PROFESSIONAL ASSOCIATIONS AND SERVICE Reviewer for Geology, Journal of Geophysical Research, Icarus, Planetary & Space Science External grant reviewer for NASA Mars Fundamental Research Program Planetary Geomorphology Image of the Month contributor (10/2008, 5/2010) American Geophysical Union (since 2001) Geological Society of America (since 2003) British Society for Geomorphology (since 2005) Association Internationale pour l’Etude des Argiles (since 2009) NASA Academy Alumni Association (since 2002) Executive Selection Board, 2005-9; Soffen Travel Grant Committee, 2006-9; Phone Interviewer, 2004, 2007, 2010

UNIVERSITY COMMITTEES AND SERVICE University Resources Committee (sets annual operating budget), graduate student representative, Brown University (2008-2010) Rhodes/Marshall Scholarship Nominating Committee, Brown University (2006-2010) Geoclub (geoscience graduate student group), Treasurer (2007); Rep. to Grad Student Council (2008) Rhodes Scholar Southern Africa Forum, executive committee member (2005-6) Committee on Environmental Quality, Washington University, student rep. and co-chair (2002-4) Student Union, Washington University, Senator, Academic Affairs committee co-chair (2001-4)

TECHNICAL SKILLS Computer: Proficient in ENVI, IDL programming, ArcGIS, Adobe Illustrator and Photoshop, MS Word, MS Excel, MS PowerPoint, MS Publisher, CorelDraw. Comfortable in Unix or Windows OS

Field techniques: Proficient in use of portable ASD spectrometer, thermal emission spectrometer, thermal camera, GPS. Experienced in GTS/theodolite surveying (both land and shallow water), transect and grid ecological surveys. Skilled in sediment core acquisition and marine grab sample acquisition and vibracoring.

Laboratory techniques: Experience with flux fusion sample preparation, ICP elemental analysis, XRD analysis of clays, laboratory sample preparation and data analysis for VNIR spectroscopic measurements

SCUBA: PADI Advanced Open Water Diver

LANGUAGE SKILLS English (native), Spanish (highly proficient), French (beginner)

ix Acknowledgements

It takes a village to complete a Ph.D.

First, friends--essential for getting through grad school. Wes, Noah, Rachel, Carolyn, Billy, Linda, Ashley and Gillian all welcomed me into what grad student life should be: regular drinks at the GCB after work over a wine and cheese board, the glory of “Rolling Stones” IM sports, Greenway, and RIPUL summer frisbee. Gillian, thanks for helping me find my way to PVD and showing me some of the gems (I still take nearly everyone to Modern Diner). Joe, thanks for being there to geek out with over food – from latkes to lagers and chorizo to purim. You’ve attained a sense of joie de vivre that I aspire to. To both Joe and Sam, for raising the level of political discourse (which I love)--I’ve enjoyed holding the tenuous middle in many debates. Peter, for introducing me to the great state of Maine and the best things about New England: skiing, weekends by the lake, lobstah bakes, and guitar duets. These days were the best part of my time here. And I will always value your support and the countless times you’ve been there to chat after a tough day at the office. To Debra, this year has been so much more fun, sharing more of it with you. A big thank you for keeping me sane during dissertation writing. The same to Ailish, best officemate ever. Uly, I’ll circumnavigate any island with you, anytime. Laura, Leah, Angela, Bronwen – let’s keep doing fun things together. Brendan, Kerri, Mark, Kate, David, and the rest of the gang: you’re awesome. Enough said. It makes me sound ancient when I say “the younger generation”, but thanks to the younger generation of grad students for carrying the banner: Jenny, Brandon, Janette, Liz, Xi, Scott, Mary, Diane, Colin—I wish I were around longer to get to know you better.

Thanks to Nancy, Celeste, and Ruth who have helped me in innumerable ways with things large and small. Lynn Carlson, you’ve helped me so much on my research. Thanks to Jay and Caleb for being my computer gurus of last resort -- much appreciated.

An enormous thank you to the CRISM group, Brown-side: J.R. Skok, Mark Salvatore, Janette Wilson, Mathieu Vincendon, and Bill Fripp. More people = more fun science and it’s been wonderful to be a part of the team with you. Too much cannot be said about those who’ve since moved on to bigger and better things. Shannon Pelkey and Leah Roach were amazingly supportive when I began dealing with CRISM data. I am grateful to have to sat next to Leah for three years. Science is so much better when you can talk about it, and Leah has an amazing ability to cut right to the heart of any question. And she could not possibly have been a nicer, more patient person when I interrupted her dissertation writing for the umpteenth time with an ArcMap question! I understand now what it took for you just to smile and answer :).

Leaving Oxford and turning down some other west coast institution to come to Brown was an incredibly tough decision. I wasn’t sure I had made the right call. But I came here because I wanted to work with Jack. I was excited by the emerging mineralogic story you told me and continually am amazed by the breadth of work that you do. May I be like you. Thank you for convincing me that I should check out phyllosilicates (I seem to recall being strongly pro- salts instead). You were right ;). But you do still owe me a bottle of wine. I recall a bet: something about finding no new mineral classes using CRISM?! Seriously, though, I am so grateful for the conference opportunities, supporting me to go off to work in Paris and at Indiana, for Iceland – who knew 1 day fighting jet lag at Hvalfjordur would lead to so much?--and for letting me be largely independent to pick and choose my opportunities, but being there when I needed you. This has been fun. I’m excited (maybe ExCITEd?) about the future.

Thanks to my professors. To Carle, I want to run a mission like you and also be an amazing classroom teacher. I hope I can fill those shoes one day. To Pete, I first met you at Wash U, and your enthusiasm helped convince me Brown was the place to be. To Mike, you’ve been an invaluable source of advice on Mars research and on post-PhD life. To Jim, you have helped link me to work in Russia and across the world, and thank you for continually exposing all of LF to the breadth of planetary sciences and its cast of characters. Greg, thanks for balancing out my committee of planetary scientists and making me think about my work in new ways. Reid and Alberto, I will forever have an appreciation of stable isotopes and thermodynamics because of you, but thanks also for taking the time to be mentors outside of the classroom.

Then there’s the whole CRISM team. Wow, it’s been exciting since October 2006! How many new-for-Mars minerals still out there? Thanks to Roger, Gregg, Dick, Janice, Wendy: you paved the way for me in VNIR spectroscopy—something about shoulders and giants would be appropriate, I think. Thanks to Mario, Ralph, Nancy, Sandra, Kim, Kim, Frank, and Debra for keeping the mission running and keeping the spectral discussions real. And I owe a huge debt of gratitude to Scott Murchie for building this incredible instrument, amassing this great collection of scientists, and being so incredibly supportive and generous to all us grad students on the team. Thank you. May all missions be the same.

Dave DesMarais, Steve Squyres you’ve been great mentors—fabulous to have you in my corner for the long-term. James Wray, it’s been wonderful to work with you—yay for CRISM-HiRISE collaboration! Dawn Cardace, may there be many future collaborations to come! And may I have half your enthusiasm. Dave Bish thank you for being so generous with your time and teaching this naïve spectroscopist the in’s and out’s of how a real clay mineralogist does XRD. Thanks to Dave Murray for patience helping me “play real scientist” complete with goggles, lab coat, and a fiery 1000°C furnace. And a big thank you to Taki for running the RELAB, including my seemingly endless queue of samples, and also being a valuable scientific sounding board for musings on radiative transfer theory.

And on the European side. Jean-Pierre, Francois, Michel, Brigitte, Yves, John, Joe, Cedric: thank you for welcoming me into the IAS fold. I’m looking forward to the next chapter.

Because this is the culmination of 22 years of education, not just four, thanks also to Carolyn Nicholson (7th & 8th grade) for teaching me how to write and Melinda Hilsenbeck and Sue Newman (high school) for teaching me how to think critically, think deeply, and to embrace the vagaries of life with an inquiring mind, an international outlook, and a sense of adventure.

I wouldn’t be here without Ray Arvidson, who somehow saw that the 18 year-old prospective ugrad in his office, who couldn’t decide between ecology and IR, politics or astronomy, earth sciences or law, might find a way to somehow meld all of those at Washington University. You trained me how to be a precise and careful scientist and also how to learn independently and teach myself. And you truly opened up the world for me, bringing a freshman along to rover field trials at JPL and then trusting that a 22-year old could run mission critical rover ops. It hooked me on planetary science for life.

Thanks to John Boardman and my classmates from around the world in the Oxford ECM MSc 2005. Whether in the classroom, over a pint (or three), or dancing at our potlucks, you made me feel centered in the greatest community in the world. I am grateful always to Heather Viles who taught me how to work independently, write a good (or at least decent) literature review, and who showed me how to do field work well, while always laughing (usually at myself).

Thanks to close friends from other chapters in life: Darryl Hill, Dan Cavanagh, Anjali Nayar, Robin Rotman, Jen Griffes, and Heather Long. You’ve had to listen to the ups and downs of this PhD, and I’m grateful for you always. Rose Kontak, you have made my last semester here better than I ever hoped. And, of course, thanks to my family. To my Dad, well, some of your academic personality must have rubbed off on me—there’s soon to be another Dr. Ehlmann in the house! Thank you for always supporting me. I love you. Thanks to Barbara for being so supportive and providing such wise and grounded advice when I needed to be pulled back to seeing life’s big picture. And thanks to my brother Jonathan and Aunt Sue for sharing this academic journey with me. And, finally, to my mother, who isn’t here anymore but who I know would be incredibly proud.

Geosciences has been such an exciting field of PhD study that has let me travel the world, peer at new worlds, and experience the excitement of discovery. Space may be the final frontier, but I’ve learned that the key is finding good company for the journey. Thanks for sticking around for the ride. Ad astrum.

Bethany Ehlmann 15 April 2010 Providence, RI xii

TABLE OF CONTENTS

Chapter 1: Introduction……………………………………………………………………1

Chapter 2: Orbital Identification of Carbonate Bearing Rocks on Mars……………...... 25

Chapter 3: Identification of Hydrated Silicate Minerals on Mars: Geologic Context Near the Nili Fossae & Implications for Aqueous Alteration………..47

Chapter 4: Geologic Setting of Serpentine-Bearing Rocks on Mars…………………...137

Chapter 5: Evidence for Low-Grade Metamorphism, Diagenesis, and Hydrothermal Alteration on Mars from Phyllosilicate Mineral Assemblages…………155

Chapter 6: Weathering and Hydrothermal Alteration of Basalts in Iceland: Mineralogy from VNIR, TIR, XRD and Lessons for Linking Orbital and Surface Data on Mars…………………………………………………………………...204

Chapter 7: Modeling Modal Mineralogy of Laboratory Mixtures of Nontronite and Mafic Minerals using Visible Near-Infrared Spectral Data: An accuracy assessment………………………………………….………………251

Chapter 8: Conclusions…………………………………………………………………292

xiii 1

Chapter 1:

INTRODUCTION

1 2

This work investigates the nature and timing of aqueous processes on ancient

Mars, focusing on the earliest epoch of Mars history, the Noachian (>3.7 Ga) and strata bearing aqueous alteration minerals. This dissertation consists of two parts. The first (Ch.

2-5) is Mars data-focused. Orbital visible/near-infrared (VNIR) imaging spectroscopic data from the Compact Reconnaissance Imaging Spectrometer for Mars (CRISM) is used to identify key minerals and characteristic mineral assemblages. By coupling mineral identifications with high resolution morphology to better assess stratigraphic relationships, the geochemistry and setting of aqueous alteration (hydrothermal, lacustrine, or pedogenic) is established or constrained. The second part of this dissertation

(Ch. 6-7) focuses on the limitations and synergies of mineralogic determinations derived from VNIR spectroscopy versus other techniques. Rocks from neutral to alkaline weathering and hydrothermal systems in Iceland were characterized with VNIR, thermal infrared (TIR), XRD, and geochemical analysis to simulate orbital and rover based analyses. VNIR spectra from manmade mixtures of smectite clays with mafic rocks were prepared to test the efficacy of Hapke and Shkuratov radiative transfer theories in estimating modal mineralogies remotely.

The broader goals sought by this work are three-fold: (1) a refined understanding of what is knowable about past aqueous conditions based on remote VNIR analysis of altered surfaces; (2) a greater understanding of the changing habitability potential of Mars through time and, more broadly, (3) an increased understanding of aqueous environments during the first billion years of the evolution of a terrestrial planet. Below I briefly review the motivation for this research and place it in the context of previous studies.

Motivation: the earliest aqueous environments on terrestrial planets

For the solar system’s terrestrial planets, the first billion years was a critical time.

Planetary crusts initially formed and cooled. Planetary surfaces were bombarded by impacts, probably including a late heavy bombardment at ~3.8 Gyr (Ryder et al., 1990;

Gomes et al., 2005). The first planetary atmospheres were established and cooled so as to allow liquid water to condense. On Earth, organisms first evolved in this time period

(Figure 1). Yet, after the first billion years, the trajectories of Earth, Mars, and Venus diverged starkly in volcanic/tectonic history, atmospheric composition and evolution, and the fate of liquid water in the surface environment. Liquid water serves as a solvent and is probably the most critical ingredient for life as we currently understand it. Hence, tracing the fate of water though time is a means for understanding the evolution of planetary habitability and for assessing the likelihood that planetary bodies other than Earth could have supported life.

Some of the key geologic questions for origins of life studies on Earth include:

When was liquid water first stable? When did Earth become habitable/inhabited? What was the nature of the first habitat(s)? The oldest, albeit controversial, fossil evidence for life on Earth dates from 3.4 Gyr (Schopf, 2006). The oldest surviving terrestrial water- lain sedimentary rocks date from ~3.8 Gyr, and zircon grains may preserve evidence for water-rock interaction as early as 4.4 Gyr (Valley et al., 2002; Moorbath, 2005).

However, the record of early Earth’s habitability and the nature of its first aqueous environments has largely been destroyed due to the recycling and deformation of crustal rocks via plate tectonics. Far less than 1% of the crust from Earth’s first billion years, the

Hadean and early Archaean, is preserved.

In contrast to Earth, Mars has substantial intact remnants of crust from this time period exposed at the surface. Approximately 40% of Mars surface is Noachian (>3.7

Gyr) and 30% is from the Hesperian (Scott & Tanaka, 1986; Greeley & Guest, 1987;

Tanaka & Scott, 1987). Mars’ Noachian epoch is comparable to the Hadean and early

Archaean eras on Earth, and includes late heavy bombardment, the earliest fluvial valley networks, and the formation of alteration minerals (Figure 1). The preservation of this

Mars record provides a fabulous opportunity to understand the fate of liquid water during the early evolution of a terrestrial planet and the consequences for its geologic and atmospheric evolution.

The climatic and geologic history of Mars is of particular relevance to understanding the long-term sustenance (or not) of environments conducive to life, viewed through the lens of the availability of water. On Mars today (Patm= ~6 mbar,

Tsurface=140-280 K), liquid water is unstable to sublimation as well as evaporation over most of the surface. Nonetheless, there is substantial geomorphologic and mineralogic evidence indicating liquid water was an active and important geologic agent on ancient

Mars. Branching valley networks indicate well-developed drainage systems (Irwin et al.,

2005; Fassett and Head, 2008), and thousands of kilometer long outflow channels empty into the northern lowlands of Mars (e.g. Carr, 1996). Mineralogic evidence, including the presence of hydrated silicate minerals and sulfates in lithologic units from the Noachian and Hesperian, respectively, indicate past aqueous chemical alteration (Bibring et el.,

2006). What was nature of the first aqueous environment(s) on ancient Mars? How and why did these environments change through time and vanish? Were these earliest environments habitable? These are the essential motivating questions of this dissertation. 5

Water on Mars: A morphologic view

The morphologic effects of water in shaping the surface landscape of ancient

Mars are pronounced. Valley networks and thousands of kilometer outflow channels were first recognized in Mariner 9 images (see for review Squyres, 1984). Subsequent observations have refined knowledge of the timing and nature of fluvial activity. Some have posited the existence of a northern ocean (Parker et al. 1993) but certainly there is evidence for standing bodies of water in open and closed basin lakes (Cabrol & Grin,

1999; Fassett & Head 2008). Well-developed valley networks (Malin and Edgett, 2003;

Mangold et al, 2004; Irwin et al., 2005), evidence for groundwater sapping (Grant, 2000), and large-scale denudation of the landscape by fluvial processes (Craddock & Maxwell,

1993) all point to a once active hydrologic cycle. Rapid release of large volumes of water formed thousands of kilometer long outflow channels (Baker and Milton, 1974).

A long-lived surface hydrologic cycle, driven by rainfall or snowmelt, could explain well-developed valley networks, putative paleolakes, filled craters and intensely degraded Noachian-aged surfaces. However, most valley network activity appears to be confined to a relatively narrow time interval in the Late Noachian/Early Hesperian

(Craddock & Maxwell, 1993; Irwin et al., 2005; Fassett & Head, 2008) (Figure 1). Large magnitude floods and activity within the catastrophic outflow channels mostly occurred later in the Hesperian (Tanaka, 1997). In the last three billion years, liquid water appears to have played a negligible role in shaping the landscapes of Mars. Recent observations of gullies and near-surface polar ice (Malin et al., 2006; Byrne et al., 2009; Smith et al.,

2009) suggest that liquid water may continue to episodically play a local role in chemical alteration and physical erosion, although its participation is likely sporadic, and brief, driven by the influence of variations in orbit and obliquity (Laskar et al., 2002). 6

Initial studies of mineralogy: A basaltic planet with hints of alteration

In contrast to the morphologic study of Mars, mineralogic study of Mars rocks and soils has been comparatively slower to progress. The Viking lander missions provided geochemical data (e.g. Toulmin et al., 1977) indicating that the soils of Mars were mafic in composition but likely with an additional salt component, perhaps added via reaction with volcanic gases (Baird & Clark, 1981). The mafic component could be primary or secondary (e.g. smectite clays); data are equivocal concerning the mineralogy

(Arvidson et al., 1989). Indeed, the first landed mission to Mars with the ability to determine the modal mineralogy of a sample will be the 2011 Mars Science Laboratory, with its ChemMin instrument, a combined XRD/XRF. Most of knowledge of the mineralogy of Mars’ rocks and soils is derived from remote telescopic and orbital observations (Table 1) and recent in-situ studies of the MER rovers.

Pyroxene was discovered on Mars using telescopic reflectance spectroscopy

(Adams and McCord, 1969; Singer et al., 1979), and both low- and high-calcium pyroxene were found with later analyses of Phobos-2 ISM datasets (Mustard & Sunshine,

1995). An absorption edge in the visible and a weak band near 0.86 m were used to infer the presence of Fe3+ and oxidized products, presumably resulting from weathering of primary mafic minerals, especially in Martian bright regions. Palagonite (Evans and

Adams, 1980) and nanocrystalline hematite (Morris et al., 1989) were put forth to explain the data. Weathering and the existence of hydrated minerals were further indicated by the presence of a band near 3 m (Sinton, 1967; Houck et al., 1973), and infrared thermal emission spectra of Mars from the Mariner spacecraft (2.0-14.0 m) were interpreted to be indicative of hydrated mineral phases (Pimentel et al., 1974). 7

Throughout the 1970s to 1990s various alteration minerals were put forth to explain the spectral properties of various telescopic datasets. An identification of carbonate was made (Pollack et al., 1990) but was not confirmed by others (Blaney &

McCord, 1990). Hydrous carbonates (Calvin et al., 1994) were put forward to explain the apparent presence of some diagnostic bands but absence of others. The aluminium smectite clay montmorillonite (Hunt et al., 1973), the bicarbonate bisulfate scapolite

(Clark et al., 1990), and sulfates (Blaney and McCord, 1995, Pollack et al., 1990) were also proposed. However, these identifications were challenging because observations were hindered by low spatial resolution and the difficulties of removal of both Earth and

Mars atmospheric effects. Few accepted the detections as unambiguous.

The dearth of mineralogic evidence of past water was paradoxically difficult to reconcile with the abundant morphologic evidence of its role on early Mars.

Nevertheless, at the time numerous lines of research continued to point to the fact that

Mars might be expected to have clays, sulfates, and carbonates, perhaps at great abundances. Models for evaporite deposits produced from standing bodies of water in equilibrium with the Mars atmosphere predicted carbonates, sulfates, and halites in paleolake basins (Catling, 1999). Models for weathering of mafic silicates predicted hydrated silicates such as smectites, serpentines, oxides and carbonates (Gooding, 1978).

A growing body of evidence suggested the SNC meteorites came from Mars, and these had evidence for clays, salts, and products of aqueous alteration (e.g. Gooding et al.,

1992).

However, over subsequent decades even higher resolution instruments (Table 1), continued to show scant evidence for substantial aqueous chemical alteration. Global mapping in the thermal infrared with TES and deconvolution of TIR spectra permitted 8 excellent definition of the distribution and abundance of major silicate minerals: olivine, pyroxenes, and feldspar allowing distinctive mafic terrains to be discriminated

(Christensen et al., 2000a; Bandfield, 2002; Rogers & Christensen, 2007). The evidence for water-related minerals commensurate with the geomorphic evidence was largely absent, however. A thousands of square kilometer unit enriched in gray crystalline hematite was found in the equatorial region of Terra Meridiani, and probable formation mechanisms all suggested liquid water (Christensen et al., 2000b). However, there was no evidence for clay minerals or carbonates or sulphate evaporite deposits. A thin Si-coating or alteration rind was proposed to account for the spectrally distinct signature of the northern plains (Wyatt & McSween, 2002). There was evidence for very small quantities of carbonate (Banfield et al., 2003) and zeolite (Ruff, 2004) in the dust at <5% levels.

Collectively, the data seemed to indicate that Mars had, in fact, likely been cold and dry throughout most of its history (e.g. Bandfield et al., 2003). Although morphologic indications of water were clear, they represented intense but geologically brief periods of time in which waters transported materials but had insufficient time to reach with mafic materials to substantially alter their mineralogy (Christensen, 2005).

New mineralogic evidence of water on early Mars

Interest in ancient Mars as a potentially watery and habitable environment was reinvigorated by the definitive discovery in 2004 of hydrated minerals in rock units.

Using visible/near infrared imaging spectroscopy, the Observatoire pour la Minéralogie l’Eau les Glaces et l’Activitié (OMEGA) on the European Space Agency’s (ESA) Mars

Express mission, discovered two classes of minerals in bedrock units: phyllosilicates and sulfates (Bibring et al., 2005; Gendrin et al., 2005; Poulet et al., 2005). The 9 phyllosilicates are predominantly iron magnesium smectites (e.g. nontronite and saponite) with a few instances of the Al-smectite montmorillonite. Phyllosilicates are found only in Noachian terrains and are exposed by erosion or impacts (Poulet et al.,

2005). Sulfates found by OMEGA include gypsum in north circumpolar dunes and variously hydrated sulphates (e.g. kieserite, polyhydrated sulfates) in rock units in and around Valles Marineris (Gendrin et al., 2005) that are also associated with hematite

(Bibring et al., 2007). The sulfate-bearing rocks are located in Hesperian (~3.7 to 3.1 Ga) aged terrains. Both phyllosilicates and sulfates are notably absent in younger Amazonian

(<3.1 Ga) rock units.

Study of the geologic setting of these hydrated minerals, their relative stratigraphy, and geochemical modelling led to the hypothesis that phyllosilicates and sulfates are marker minerals for distinctive periods in the evolution of Mars’ aqueous environments (Bibring et al., 2006). In this paradigm for Mars evolution, the first several hundred million years following formation of the crust was a phyllosilicate-forming era in which liquid water of neutral to alkaline pH was present in sufficient quantities to leach volcanic rocks and alter primary minerals. This era was followed, starting ~3.7 billion years ago, by a more acidic and water-limited era of evaporative sulphate formation. Following a period of global change involving enhanced volcanic activity and loss of atmosphere, increased pSO2 and decreased pH2O created acidic evaporative conditions favorable to sulfate formation (Bibring et al., 2006; Chevrier et al., 2007).

Concurrent with the OMEGA investigation, the Opportunity rover discovered mineralogic and sedimentologic evidence consistent with evaporative sabkha type environments at Hesperian Meridiani Planum (Squyres et al., 2004; Grotzinger et al.,

2005). The conditions at this site, with episodic surface waters and diagenesis (McLennan 10 et al., 2005), are perhaps, exemplary of drying global conditions during this time period inferred from orbital data. For the past 2-3 billion years, liquid water has not been stable on the surface of Mars, and alteration has been anhydrous leading to abundant ferric oxides.

With the addition of a second higher resolution VNIR imaging spectrometer on the Mars Reconnaissance Orbiter (MRO), CRISM, in 2006 (Table 1), it has become clear that phyllosilicate alteration minerals in the southern highlands are globally widespread

(Mustard et al., 2008). Phyllosilicates have now been identified in hundreds of locations.

The number of alteration minerals definitively detected in rocks of the ancient crust of

Mars has also only continued to grow (Figure 2). Carbonates, zeolites, other hydrated silicates and sulfates have been found, pointing to a greater diversity of aqueous environments on early Mars than had been previously known (Mustard et al., 2008;

Murchie et al., 2009), as will be further discussed herein.

Key Questions, Approach, & Dissertation Summary

This dissertation work began with first analyses of the new data coming back from CRISM. As more exposures of hydrated silicate minerals have been found and more diverse mineral classes discovered, the paradigm (Bibring et al., 2006) of a general transition from phyllosilicates in Noachian-aged strata to exclusively sulfates in

Hesperian-aged strata has largely held. Numerous questions, however, remain concerning environmental conditions during the first, clay forming era. Was phyllosilicate formation a global or a local phenomenon? Are phyllosilicates related to surface waters and valley formation or to other processes during the Noachian, e.g. crustal cooling, late heavy bombardment? Did a single homogenous alteration process operate or are there multiple 11 formation mechanisms to explain the minerals discovered? Were the settings in which phyllosilicates formed potentially habitable environments? Are all instances of phyllosilicates restricted to the Noachian? What was the nature of the transition to the sulfate-forming epoch?

This dissertation treats many of these questions using new CRISM data and analyses of terrestrial materials. It is comprised of two parts. Part 1 (Chapters 2-5) focuses on analysis of VNIR data from Mars, identifying new alteration minerals and placing them in stratigraphic context to determine likely formation environments. Part 2

(Chapters 6-7) focuses on VNIR imaging spectroscopy as a technique, in particular assessing the limits of its accuracy for qualitative and quantitative mineralogic analyses.

The chapters are summarized briefly below.

Part 1: Mars aqueous mineralogy and stratigraphy

Following data calibration to radiance and correction for atmospheric and photometric effects, for each pixel in a CRISM image (best resolution, 18m/pixel) minerals can be identified from a 544 channel VNIR spectrum of the surface. This allows mapping of minerals spatially and correlation with rock units and geomorphic features like layered units, ridges, channels, and craters. This context is essential to moving from fundamentally observations of chemistry to observations of geology, which allow the chemistry to be time-ordered and used to determine the nature of ancient environments.

Chapter 2 reports the discovery, geologic setting, and implications of carbonate-

bearing rock units on Mars. Along with several hydrated silicate minerals

discovered during CRISM primary science operations (Mustard et al., 2008;

Milliken et al., 2008), carbonate provides evidence for multiple distinctive

alteration environments on early Mars. In this case, large exposures of carbonate- 12 bearing rocks formed in neutral- to alkaline-pH environments, probably generated by buffering waters with regionally extensive ultramafic rocks circumferential to the Isidis basin (Ehlmann et al., 2008).

Chapter 3 follows up on the discoveries in Chapter 2 and on prior work (Mustard et al., 2007; Mangold et al., 2007) mapping the relationships between alteration mineral-bearing units west of the Isidis basin in and around the Nili Fossae. A careful, detailed rationale is also provided for the identification of each new hydrated silicate discovered by CRISM: chlorite, serpentine, kaolinite, illlite/muscovite, hydrated silica, and analcime. The mineralogic diversity occurs in several geographically discrete “provinces” of alteration with distinctive assemblages of alteration minerals. Multiple episodes of aqueous alteration including both near-surface leaching/pedogenic environments and subsurface hydrothermal environments are implied for this region of Mars (Ehlmann et al.,

2009).

Chapter 4 provides a global map of serpentine-bearing rock units on Mars and a description of serpentine’s various geologic settings. Serpentinization of ultramafic rocks occurred in the Noachian and was a past source of methane

(Ehlmann et al., 2010).

Chapter 5 synthesizes evidence for hydrothermal, low-grade metamorphic and/or diagenetic alteration on Mars from mineral assemblages including prehnite, analcime, and illite/muscovite, all newly discovered by CRISM. Terrestrial analogs and formation settings are considered (Ehlmann et al., submitted).

Part 2: VNIR spectroscopy techniques

Until rovers and landers visit some of the sites of mineralogic alteration discussed herein, orbital data remains the sole source for characterization of these sites. At present, only VNIR datasets can reliably identify hydrated minerals from orbit. However, although enormous strides can be made with orbital VNIR spectroscopy in terms of understanding past environments through aqueous mineralogy, it is necessary to fully characterize the limits as well as the benefits of the technique and to understand how orbital data relates to more familiar in-situ techniques such as XRD and elemental analyses.

Chapter 6 reports initial results from a study of weathered and hydrothermally

altered basaltic rocks from Iceland, from neutral to alkaline environments thought

to be analogous to some of those recently inferred on Mars by analyses of CRISM

data. The mineralogy of samples was determined qualitatively with VNIR and

TIR spectral characterization and compared to quantitative XRD and elemental

analyses such as will be performed by the MSL rover.

Chapter 7 reports initial results from tests of the efficacy of the Hapke (1993) and

Shkuratov (Shkuratov et al., 1999) radiative transfer models in determining the

modal mineralogy of mixtures bearing nontronite, an Fe smectite clay, and mafic

minerals using VNIR spectral data of prepared mixtures. A goal of radiative

transfer modelling is to move from qualitative VNIR mineral identifications

toward quantititative modal mineralogy. A necessary first step is laboratory

validation of existing radiative transfer models and improvements to those

models. This has been done for anhydrous mixtures (Mustard & Pieters, 1989; 14

Poulet & Erard, 2004) but has not been well-tested for mixtures with hydrous

minerals.

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Table 1. A summary of spectroscopic observations of Mars’ surface Time Period Data Source Wavelength Spatial Resolution Notes 1940-1990s Earth-based telescopic various 100s km scale (and continuing) observations wavelengths 1969 NASA Mariner 6/7 Infrared 1.9-3.7 m, Large spatial scale 240 spectra Spectrometer (IRS) 3.0-6.0 m acquired 1989 Soviet Phobos, French Imaging 0.7-3.2 m 22 km/pixel 40,000 Spectrometer for Mars (ISM) spectra acq. 1996 Thermal Emission Spectrometer 6-50 m 3x5 km/ spot size (TES) Dec. 2003 ESA Mars Express OMEGA 0.4-5.0 m 300 m- 5 km/pixel First NIR imaging spectrometer Oct. 2006 NASA Mars Reconnaissance 0.4-4.0 m 18 m - 200 m/pixel Orbiter CRISM

Figure 1. Global geologic timeline, focusing on key events and periods in early Earth and Mars history. 1) Nyquist et al., 2001 (2) Borg et al., 1999 (3) Valley et al., 2005 and references therein (4) Schopf, 2006 (5) e.g. Farquhar et al., 2000. 24

Figure 2. Minerals detected on Mars from remote spectroscopic observations. Since 2005, a remarkable expansion in the known mineralogic diversity of the Martian surface has been brought about with orbital visible/near-infrared spectroscopy. (synthesized from Soderblom, 1992; Calvin & Bell, 2008; Murchie et al., 2009).

Chapter 2:

ORBITAL IDENTIFICATION OF CARBONATE-BEARING

ROCKS ON MARS

Bethany L. Ehlmann1, John F. Mustard1, Scott L. Murchie2, Francois Poulet3, Janice L. Bishop4, Adrian J. Brown4, Wendy M. Calvin5, Roger N. Clark6, David J. Des Marais7, Ralph E. Milliken8, Leah H. Roach1, Ted L. Roush7, Gregg A. Swayze6, James J. Wray9

1. Department of Geological Sciences, Brown University 2. Johns Hopkins University/Applied Physics Laboratory 3. Institut d'Astrophysique Spatiale, Université Paris Sud 11 4. SETI Institute and NASA Ames Research Center 5. Department of Geological Sciences and Engineering, University of Nevada 6. US Geological Survey, Denver 7. National Aeronautics and Space Administration, Ames Research Center, Mountain View, California 8. Jet Propulsion Laboratory, California Institute of Technology 9. Department of Astronomy, Cornell University

Published in:

Science

December 19, 2008

Abstract

Geochemical models for Mars predict carbonate formation during aqueous alteration.

Carbonate-bearing rocks had not previously been detected on Mars’ surface, but Mars

Reconnaissance Orbiter mapping reveals a regional rock layer with near infrared-spectral characteristics consistent with the presence of magnesium carbonate in the Nili Fossae region. The carbonate is closely associated with both phyllosilicate-bearing and olivine- rich rock units and likely formed during the Noachian or early Hesperian from alteration of olivine by either hydrothermal fluids or near-surface water. The presence of carbonate as well as accompanying clays suggests that waters were neutral to alkaline at the time of its formation and that that acidic weathering, proposed to be characteristic of Hesperian

Mars, did not destroy these carbonates and thus did not dominate all aqueous environments.

Main Text

Although telescopic measurements hinted at the presence of carbonate on Mars (1-3), subsequent orbiting and landed instruments found no large scale or massive carbonate- bearing rocks (4, 5). Carbonate in veins within Martian meteorites (6) and possibly at

<5% abundance in Mars dust (1, 4) indicate it is present as a minor phase. The lack of carbonate-bearing rock outcrops is puzzling in light of evidence for surface water and aqueous alteration, which produced sulfate and phyllosilicate minerals (5, 7). Carbonate is an expected weathering product of water and basalt in an atmosphere with CO2 (8, 9), and large scale deposits, which might serve as a reservoir for atmospheric CO2, were predicted for Mars (10). Lack of carbonate among identified alteration minerals has compelled suggestions that either (i) a warmer, wetter early Mars was sustained by

greenhouse gases other than CO2 (11, 12), (ii) liquid water on Mars’ surface in contact with its CO2 atmosphere was not present for long enough to form substantial carbonate

(13) (thus implying minerals such as phyllosilicates must have formed in the subsurface), or (iii) formation of carbonate deposits was inhibited or all such deposits were destroyed by acidic aqueous activity (14, 15) or by decomposition (16). Here we report the detection of carbonate in a regional-scale rock unit by MRO’s Compact Reconnaissance

Imaging Spectrometer for Mars (CRISM) and discuss the implications for the climate and habitability of early Mars.

In targeted mode CRISM acquires hyperspectral images from 0.4-4.0 μm in 544 channels at a spatial resolution of 18 m/pixel (17). In addition to diverse hydrated silicates (18), CRISM identified a distinct, mappable spectral class of hydrated material in the Nili Fossae region and two nearby areas (Fig. 1) for which a match to known mineral reflectance spectra was not initially evident (19). This spectral class has a 1.9 μm combination overtone from structural H2O and also characteristic absorptions at 2.3 and

2.5 μm and a broad 1 μm absorption (Fig 2). Similar spectra have been obtained with the

Observatoire pour la Minéralogie, l’Eau, les Glaces, et l’Activité (OMEGA) imaging spectrometer on Mars Express (20).

Magnesium carbonate is the best candidate to explain the distinctive features of this spectral class. Paired absorptions at 2.3 and 2.5 μm are overtones and combination tones of C-O stretching fundamentals in the mid-infrared (21). The wavelength of their minima identifies the major metal cation in the carbonate (Fig. 3) (21, 22). Anhydrous carbonates with mostly Mg exhibit minima at shorter wavelengths (2.30, 2.50) than those with mostly Ca (2.34, 2.54) and Fe (2.33, 2.53) (22) and match the spectral class identified by CRISM (Fig. 3). We know of no other mineral spectrum that has all of the

28 properties of this class in terms of band position and width in the 2.0 to 2.6 μm spectral region (Fig. 2B-C). The overall spectral shape, position, and relative strengths of the 2.3 and 2.5 μm absorption bands are consistent among CRISM spectra (Fig. 2B, 3), suggesting that the distinctive spectral class is generated by the presence of a single phase rather than a mixture of many alteration minerals (23). In a mixture, the relative strengths of individual bands would be expected to vary with variation in the relative abundances of the mineral components. In some CRISM spectra of what we infer to be carbonate- bearing materials (Fig. 2B), intimate or spatial mixing is indicted by a wavelength shift and narrowing at 2.3 μm and broadening at 2.5 μm that is accompanied by appearance of a weak band at 2.4. These collectively indicate the presence also of iron-magnesium smectite (e.g. nontronite, Fig. 2C-orange), which has been previously identified in the region (18-20, 24).

Although the distinctive CRISM spectroscopic signature was recognized in earlier

OMEGA (20) and CRISM (19) observations, a carbonate mineral identification was rejected because the data lacked the strong 3.4 and 3.9 μm overtone absorptions seen in some laboratory data of calcite, other anhydrous carbonates (green, Fig. 2E), and their mixtures (25). However, we found the highly correlated 2.3 and 2.5 μm bands in many

CRISM observations and thus re-examined the 3-4 μm region in both CRISM and laboratory data. Absorptions at 3.45 and 3.9 μm in the CRISM spectra are present in terrains with the 2.3/2.5 μm absorptions yet not in terrains lacking those absorptions (Fig.

2D; Supplementary Online Material) but are quite subtle and only become apparent after averaging spectra from hundreds of pixels. Laboratory data show that absorptions from 3-

4 μm in carbonate are not always strong (e.g. purple and blue, Fig. 2E). The presence of water, coatings, or additional minerals can reduce or eliminate these features. In the

29 putative carbonate-bearing spectral class, the presence of a water-bearing phase(s) is indicated by 1.9 μm (Fig. 2B) and a deep 3.0 μm absorption (Fig. 2D). Hydrous carbonates—carbonates whose structures incorporate water—frequently have no 3.4 or

3.9 μm bands (2, 26) and are a kinetically favored low-temperature alteration product from solutions with Mg and CO3 (9, 27, 28). Strong overtones and fundamentals of water and OH near 3 μm in hydrated phases, e.g. hydrous carbonates or clays, mixed with carbonate can subdue the 3-4 μm carbonate absorptions (blue, Figure 2E; magnesite+hydromagnesite in 21, 29). Additionally, remote detections in the 3-4 μm region are complicated by a thermal emission contribution which reduces band strength

(30, 31) and also by instrument effects. CRISM has over a factor of 4 lower signal-to- noise ratio at wavelengths >2.7 μm, and interpretation of that region is additionally complicated by uncorrected out-of-order light (17) and a likely detector artifact at 3.18

μm. The combination of the subtle absorption features at 3.4 and 3.9 μm and the strong

2.3 and 2.5 μm bands is consistent with the presence of carbonate.

The CRISM spectra also display a strong broad band near 1.1 um, which is

2+ generated by electronic transitions of Fe (32). Magnesite (MgCO3) and siderite (FeCO3) form a complete solid solution, and a strong broad electronic band centered near 1.1 μm is apparent with even <1 wt. % iron, without changing the position of the 2.3 and 2.5 μm bands (21; green, Fig. 2E). Alternatively, the strong 1.1 μm band in the putative CRISM carbonate spectra might result from small amounts of olivine, which is commonly associated with the carbonate as discussed below. Indeed, a laboratory mixture of magnesite (80 wt%), olivine (15 wt%), and the Fe-rich smectite nontronite (5 wt%) produces a spectrum similar to that observed by CRISM (blue, Fig. 2E).

Thousands of CRISM targeted images sampling Mars’ surface have been examined for this phase (33). One image in Terra Tyrrhena and two in Libya Montes contain small exposures of carbonate-bearing rocks but the largest and most clearly defined exposures are in the Nili Fossae region, found to-date in 24 CRISM targeted images (Fig. 1). The magnesium carbonate is present in relatively bright rock units exposed over <10 km2, which allows detection by OMEGA and CRISM but probably precludes definitive detection by the Thermal Emission Spectrometer (TES) with its larger spatial footprint. The carbonate-bearing materials are restricted to Noachian cratered terrain (34), and their brightness in nighttime thermal infrared images and morphology in high resolution HiRISE images indicates that they occur in lithified deposits. The carbonate is observed in eroded mesa topography around the fossae, rocks exposed on the sides of valleys in the Jezero crater watershed and elsewhere, and sedimentary rocks within Jezero crater (35).

The carbonate-bearing rocks are relatively bright toned and are commonly fractured (Fig. 4). Like the regional smectite and olivine deposits in Nili Fossae (18, 36,

37), the carbonate-bearing rocks consistently lie stratigraphically beneath an unaltered mafic cap unit (Fig. 4). All scenes that exhibit carbonate also exhibit iron-magnesium smectite-bearing rock units. In many examples, the carbonate-bearing unit is clearly above the smectite-bearing unit (Fig. 4), although in some cases, the relationship is indeterminate and lateral variations between smectite and carbonate create pixels displaying mixtures of the phases. In scenes where both carbonates and aluminum phyllosilicates can be mapped clearly, the carbonate-bearing unit is always stratigraphically lower. The carbonate bearing unit appears to occupy the same

31 stratigraphic position as other nearby olivine-bearing units (18), namely beneath the mafic cap unit but above Fe/Mg smectite-bearing units.

Whereas iron magnesium smectites are found in a broad region extending westward to the Antoniadi basin (19), carbonate is restricted to the eastern portion of Nili

Fossae (Fig. 1). This area is the most olivine-rich region so far observed on Mars (38, 39) and has numerous valleys and sapping channels which indicate that extensive surface fluvial activity extended into the early Hesperian (36).

We propose two possible formation settings to explain the origin, stratigraphy, and distribution of these carbonate-bearing rocks. First, the carbonate could have formed in the subsurface by groundwater percolating through fractures in the ultra-mafic rock and altering olivine. This may have occurred at only slightly elevated temperatures, as determined by the geothermal gradient. Alternatively, hot olivine-rich rocks excavated from deep in the crust by the Isidis impact (37) or volcanic flows (39, 40) were deposited on top of water-bearing phyllosilicate rocks of the Noachian crust and may have initiated local hydrothermal alteration in a zone along the contact. The magnesite thus might occur in veined structures throughout olivine-rich rock, a relationship also observed in some

Martian meteorites (6). These ultramafic rocks might have been serpentinized; however,

CRISM has not yet conclusively identified serpentine.

An alternative explanation is that exposed olivine-rich rocks were weathered at surface ambient temperatures, perhaps during the surface fluvial activity in Nili Fossae that continued after the Isidis impact into the early Hesperian (36). The transformation of olivine rich rocks to magnesite under cold, dry conditions on Mars might resemble the weathering of olivine-rich meteorites in Antarctica (41), which produces magnesite and iron oxide mineral assemblages as rock rinds/coatings. A more water-rich surface

32 formation scenario would be that carbonate precipitated in shallow ephemeral lakes (42) from waters enriched in Mg2+ relative to other cations by percolation through ultramafic, olivine-bearing rocks. Either scenario implies that surface conditions in the Nili Fossae region were sufficiently wet to cause chemical weathering during the late Noachian or early Hesperian eras.

The Nili Fossae carbonates do not appear to have sequestered large quantities of

CO2. With the possible exception of carbonate in transported sedimentary units within

Jezero crater (35), we found no evidence of classic bedded sedimentary carbonate rocks resembling those on Earth. Instead, our results are consistent with carbonates having formed in response to unique local conditions. Although olivine is globally distributed on

Mars (43, 44), ultra-mafic rocks and their substantial interaction with water may have been necessary to generate carbonate in sufficient quantities to be detected from orbit at resolutions of tens of meters per pixel. Mg-carbonate bearing rocks found at Nili Fossae and perhaps also carbonate present at scales undetectable by CRISM, may contribute to a few percent magnesite in dust indicated by TES (4).

The existence of carbonate in rocks on Mars implies that neutral to alkaline waters existed at the time of their formation. Such conditions are consistent with those indicated by iron magnesium smectite formation during the Noachian (5, 11, 24) but contrast with the acid, low water activity conditions thought to prevail over at least some of Mars during later time periods (5, 45). The survival of the Nili Fossae carbonates indicates that they escaped destruction by exposure to acidic conditions, which would have dissolved the carbonate. Because aqueous activity in the Nili Fossae region extended into the Hesperian era (36), these carbonate-bearing rock units indicate that not all aqueous crustal environments experienced the acidic, sulfate-forming conditions

33 proposed to be characteristic of the planet during the Hesperian era, approximately 3.5 billion years ago (5, 46). Ancient Mars apparently hosted aqueous environments in a variety of geologic settings in which waters ranged from the acidic to the alkaline. Such diversity bodes well for the prospect of past habitable environments on Mars.

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Figures

Figure 1. (A) Global map of carbonate detections by CRISM (green circles) on a THEMIS day infrared image with MOLA elevation data (blue is low and red is high). (B) THEMIS day IR of the Nili Fossae region with olivine mapped from OMEGA orbits up to 4500 using the olivine parameter (43). Targeted CRISM images examined in this study are shown and outlined in yellow where Fe/Mg phyllosilicates were found and white where they were not. Green circles indicate carbonate detections. Where image footprints have significant overlap,

Figure. 2. CRISM and candidate mineral VNIR spectra. CRISM data have been photometrically and atmospherically corrected as in (18) and then processed with the noise reduction algorithm of (47). (A) CRISM reflectance spectra from hundreds of pixel regions of interest from HRL000040FF. (B) Spectral ratios, which highlight differences in composition between two units, are shown for the regions of interest in five different CRISM images (top to bottom: FRT00003FB9, FRT0000A09C, FRT000093BE, HRL000040FF, FRT0000B072) that contain the unique spectral class. The bold spectrum is the ratio from (A). (C) laboratory spectra (29; RELAB spectral database) of (top to bottom) magnesite, the hydrated Fe/Mg carbonate brugnatellite, a hydromagnesite mixture, siderite, brucite, the zeolite analcime, nontronite, serpentine, and chlorite. (D) CRISM spectral ratios over the full wavelength range, using the same denominator, for terrains from HRL000040FF inferred to be carbonate-bearing (black) and olivine-bearing (gray). (E) Laboratory spectra for fayalitic olivine, magnesite, and a mixture of magnesite (80 wt.%) and nontronite (5 wt.%), sparsely covered with medium-grained olivine sand (15 wt.%).

Figure 3. Scatter plot of continuum removed absorption band positions for anhydrous carbonates, hydrous carbonates, and other minerals with 2.3 and 2.5 μm absorptions. CRISM data that we identify as magnesite are shown as open circles within the larger dashed circle. Laboratory mineral spectra are shown from Gaffey (22; green circles), Hunt and Salisbury (21; red squares), RELAB spectra measured by E. Cloutis (orange triangles), and the USGS spectral library (29; purple circles). Band centers of CRISM spectra are known to ±0.01 μm.

Figure 4. Geomorphology and stratigraphy of the carbonate bearing units. (A) CRISM false-color composite of FRT0000B072 (R: 2.38, G: 1.80, B: 1.15 μm) where carbonate is green, olivine is yellow to brown, phyllosilicate is cyan and the mafic cap unit is purple. (B) subset of HiRISE PSP_002532_2020 from the white box in (A) showing a mafic knob overlying carbonate bearing terrain. (C) FRT000093BE with colors as in (A). (D) subset of HiRISE PSP_006778_1995 from the white box in (C) which shows the stratigraphy of carbonate-bearing units. (E) schematic stratigraphy of the mineralogic units in the Nili Fossae region (not to scale).

Supporting Online Material

Co-occurrences and spatial correlations of the 2.3, 2.5, 3.4, and 3.9 bands

In over 20 images from Nili Fossae, the material we identify as carbonate is characterized by strong 2.30 and 2.51 μm absorption bands observed in conjunction with a 1.9 μm hydration band and a broad 1.0 μm Fe2+ feature. Due to its continuum shape, the material we identify as carbonate appears characteristically bright green in false color infrared images generated with 2.38 μm as red, 1.80 μm as green, and 1.15 μm as blue

(Figure S1a). In Nili Fossae, the carbonate can also be mapped uniquely relative to other mineral-bearing units using three parameters for absorption bands: the OLINDEX parameter which is sensitive to the 1.0 μm ferrous iron band, the D2300 parameter which is sensitive to narrow vibrational absorptions near 2.3 um, and the BD2500 parameter which maps narrow vibrational absorptions near 2.5 μm (Table 1). The rationale and formulation of the first two are described in greater detail in Pelkey et al. (2007) while the latter was formulated in order to map in hyperspectral data the occurrence of the 2.5

μm absorption observed first in the Nili Fossae region (Table 1; 19). The putative carbonate bearing phase has all three absorption bands and maps white in images composed of these three parameters (Figure S1a-c; circled regions). In this parameter mapping, while a 2.3 μm absorption is frequently observed alone and serves as an indicator for regional iron/magnesium smectite-bearing rock units, a 2.5 μm absorption is never observed alone in the Eastern Nili Fossae region. It is always paired with the 2.3

μm absorption. An example from CRISM targeted image FRT00003E12 is shown in

Figure S1a-b.

The 3-5 μm spectral range is not yet commonly employed on Mars for identification of minerals due to the complexity of interpreting this region where reflectance and thermal emission both contribute; however, carbonate has overtones and combinations of fundamental bands at these wavelengths that can be diagnostic. Bands from 3-4 μm are typically strong in anhydrous carbonate and weak to absent in hydrous carbonate (Calvin et al., 1994; Cloutis et al., 2003; Jouglet et al., 2007). In addition to the apparently unique match between the shape and position of the 2.3 and 2.5 μm bands in the spectral phase we observe on Mars with those of magnesite (Fig. 2 and 3), a key factor in the identification of the phase we observe on Mars as carbonate is that it also has absorptions in the 3-4 μm region near 3.4 and 3.9 μm. The 3.4 μm band in the putative carbonate bearing phase is weak in CRISM data (its appearance at a level above the noise requires 100s of pixel averages), and it is not obviously present in OMEGA data which is acquired at ten to thirty times coarser spatial resolution. However, a parameter was constructed to map the stronger absorption at 3.9 μm in CRISM data (Table 1; Figure

S1c). (OMEGA has a strong, 10-20% non-linearity in the detector at this wavelength and cannot be employed). The BD3900 parameter appears strongest in bright-toned rock units which we infer to be olivine and carbonate bearing.

The CRISM L detector data calibration in this wavelength range is not yet fully understood, and there are some complications to mapping the 3.9 μm band, including a left to right gradient across the scene and thermal effects (and an intermittent artifact at

3.18 μm which not present in this scene). However, a true 3.9 μm band can be easily distinguished from a high BD3900 value resulting from a spectral slope by examining individual spectra. In ratioed data the putative carbonate bearing phase has absorptions at

1.0, 1.9, 2.3, 2.5, 3.4, and 3.9 while locations of other high BD3900 values do not display

43 the 2.5 μm feature and have simply strong slopes from 3 to 4 μm (Figure S1g). The 3.9

μm band is stronger in the putative carbonate bearing phase, i.e. that with the 2.3 and 2.5

μm absorptions than in other locations in the scene and is distinctly a band. In Figure

S1d-e, we use a scatterplot of all scene values for BD2500 and BD3900 to highlight regions with the highest values in both. These correspond to the putative carbonate- bearing terrains predicted from examination of the false color image and the

OLINDEX/D2300/BD2500 parameter maps (Figure S1a-b). A suite of 100’s of pixel

ROIs from putative carbonate bearing units from multiple images are shown over the full

L detector range in Figure S2. All show bands at 1.0, 1.9, 2.3, 2.5, 3.4, and 3.9 um, diagnostic of carbonate.

Additional References

S.M. Pelkey, et al., J. Geophys. Res., 112, E08S14, doi:10.1029/2006JE002831 (2007)

Table 1. Wavelengths and formulae used in CRISM parameter maps.

Parameter Formula Range OLINDEX R1695 / [0.1(R1080)+0.1(R1210)+0.4(R1330)+0.4(R1470)] -1.0 0 to 0.5 D2300 1.0 – ((CR2290+CR2320_CR2330)/(CR2120+CR2170+CR2210)) 0 to 0.07 where CR are essentially R values, continuum removed from a local slope over 1.82 to 2.53 μm BD2500 1.0 – (R2500+R2510)/(R2540+R2380) 0 to 0.05 BD3900 1.0 – AVG(R3837:R3897)/(2.0(AVG(R3597:R3663))- -0.1 to 0.05* AVG(R3364:R3430)) * BD3900 strength reflects the difference of a multi-channel average reflectance at 3.86 μm from that of the projection of a straight line through multi-channel average reflectance at 3.4 and 3.63 μm. Because of the steep upward rise of the Planck function due to surface temperature (Figure S1f), values should be very negative for no band and may be slightly negative to positive when a band is present.

45 45

Figure S1. Nili Fossae CRISM image FRT00003E12 (also discussed in 18) (A) false color composite with R: 2.38, G: 1.80, B: 1.15 um (B) parameter map with R: OLINDEX, G: D2300, B: BD2500 (C) BD3900 parameter map. Spots of putative carbonate are circled. (D) same as (A) with regions of high BD2500 and BD3900 from pixels selected in (E) overlain. (E) scatterplot of BD3900 vs. BD2500. Pixels with high values in both are selected in red (F) atmospherically corrected spectra obtained from 5x5 pixel regions with locations indicated by the boxes in (B), (G) ratios of the yellow and green spectra in (F) to the purple.

Figure S2. CRISM ratio spectra from 6 images in Nili Fossae (same as those in Figure 2b) shown over the full wavelength range of the L detector. In parentheses is the number of pixels from the putative carbonate phase used in the numerator spectrum. The number of pixels used in the denominator spectrum was always greater, and all denominator spectra had shape similar to the purple in Figure S1f above.

Chapter 3:

IDENTIFICATION OF HYDRATED SILICATE MINERALS ON MARS USING MRO-CRISM: GEOLOGIC CONTEXT NEAR NILI FOSSAE AND IMPLICATIONS FOR AQUEOUS ALTERATION

Bethany L. Ehlmann1, John F. Mustard1, Gregg A. Swayze2, Roger N. Clark2, Janice L. Bishop3, Francois Poulet4, David J. Des Marais5, Leah H. Roach1, Ralph E. Milliken6, James J. Wray7, Olivier Barnouin-Jha8, Scott L. Murchie9

1. Department of Geological Sciences, Brown University 2. U.S. Geological Survey, Denver 3. SETI Institute/ NASA Ames Research Center 4. Institut d’Astrophysique Spatiale, CNRS/Université Paris Sud 5. NASA Ames Research Center 6. Jet Propulsion Laboratory/California Institute of Technology 7. Department of Astronomy, Cornell University. 8. Johns Hopkins University Applied Physics Laboratory

Published in:

Journal of Geophysical Research

October 23, 2009

47 48

Abstract. The Noachian terrain west of the Isidis basin hosts a diverse collection of alteration minerals in rocks comprising varied geomorphic units within a 100,000 km2 region in and near the Nili Fossae. Prior investigations in this region by the Observatoire pour l’Minéralogie, l’Eau, les Glaces, et l’Activité (OMEGA) instrument on Mars

Express revealed large exposures of both mafic minerals and iron-magnesium phyllosilicates in stratigraphic context. Expanding on the discoveries of OMEGA, the

Compact Reconnaissance Imaging Spectrometer for Mars (CRISM) aboard the Mars

Reconnaissance Orbiter (MRO) has found more spatially widespread and mineralogically diverse alteration minerals than previously realized, which represent multiple aqueous environments. Using CRISM near-infrared spectral data, we detail the basis for identification of iron and magnesium smectites (including both nontronite and more Mg- rich varieties), chlorite, prehnite, serpentine, kaolinite, potassium mica (illite or muscovite), hydrated (opaline) silica, the sodium zeolite analcime, and magnesium carbonate. The detection of serpentine and analcime on Mars is reported here for the first time. We detail the geomorphic context of these minerals using data from high resolution imagers onboard MRO in conjunction with CRISM. We find that the distribution of alteration minerals is not homogeneous; rather, they occur in provinces with distinctive assemblages of alteration minerals. Key findings are (1) a distinctive stratigraphy, in and around the Nili Fossae, of kaolinite and magnesium carbonate in bedrock units always overlying Fe/Mg smectites and (2) evidence for mineral phases and assemblages indicative of low-grade metamorphic or hydrothermal aqueous alteration in cratered terrains. The alteration minerals around the Nili Fossae are more typical of those resulting 49 from neutral to alkaline conditions rather than acidic conditions, which appear to have dominated much of Mars. Moreover, the mineralogic diversity and geologic context of alteration minerals found in the region around the Nili Fossae indicates several episodes of aqueous activity in multiple distinct environments.

1.0 Introduction

Landed measurements by the Mars Exploration Rovers (MER) and orbital visible and near-infrared data from the Observatoire pour l’Minéralogie, l’Eau, les Glaces, et l’Activité (OMEGA) and the Compact Reconnaissance Imaging Spectrometer for Mars

(CRISM) have revealed a rich mineralogic record of aqueous alteration on Mars that complements pre-existing geomorphic data recording the past presence of liquid water

[Squyres et al., 2004; Bibring et al., 2005; Murchie et al., this issue]. Sulfate minerals found in northern circumpolar deposits, Vallis Marineris, the plains of Terra Meridiani, and soils around Columbia Hills in Gusev crater probably indicate precipitation of salts from acidic surface and near surface waters in the Hesperian epoch [Gendrin et al., 2005;

Bibring et al., 2006; Squyres et al., 2004; Arvidson et al., 2006; Haskin et al., 2005].

During the earlier Noachian epoch on the other hand, a different process of longer-lived water-rock interaction at neutral to alkaline pH is indicated by the presence of phyllosilicate minerals, mostly iron magnesium smectite clays [Bibring et al., 2006;

Chevrier et al., 2007]. OMEGA revealed large phyllosilicate exposures in the well- exposed Noachian terrains around Mawrth Vallis and Nili Fossae as well as in scattered outcrops in the southern highlands [Poulet et al., 2005]. Subsequent investigations at smaller spatial scales by CRISM reveal that phyllosilicates are a relatively common 50 feature of the southern highlands occurring—apparently without geographic restrictions—in association with hundreds of impact craters [Mustard et al., 2008].

Additionally, CRISM data have allowed discovery of the diversity of hydrated minerals present. In addition to previously detected smectites and chlorite [Poulet et al., 2005],

CRISM near-infrared (NIR) spectral data have been used to identify kaolinite, potassium micas such as illite or muscovite, hydrated silica, magnesium carbonate, and prehnite

[Mustard et al., 2008; Milliken et al., 2008; Ehlmann et al.. 2008a; Buczkowski et al., this issue]. Along with the detections of zeolite and serpentine reported in this study, these minerals indicate that a variety of aqueous alteration processes were probably at work on

Noachian Mars to generate such mineralogic diversity.

The Noachian epoch spans the period from the first appearance of a solid crust to approximately 3.6 Ga [Hartmann and Neukum, 2001], and the environment(s) in which aqueous alteration occurred are not yet well-defined. In comparison to the precipitation of salts from solution, phyllosilicates require a more extended period of time to form, either from topotactic alteration of mafic mineral grains or by neoformation from solution

[Eberl, 1984]. Common terrestrial processes for phyllosilicate formation include leaching and pedogenic alteration, subaqueous alteration in basins, precipitation in hydrothermal systems, and metamorphism and diagenesis. Mineral assemblage, alteration mineral abundance, and stratigraphic setting are constraints that are at least partially ascertainable from orbit and that allow distinguishing among phyllosilicate formation mechanisms.

On Mars, alteration of parent materials to phyllosilicates may have been associated with precipitation, runoff and soil formation as in the modern terrestrial hydrological cycle. Indeed, the Mawrth Vallis region shows well-developed stratigraphy 51 of Al phyllosilicates overlying Fe/Mg smectites [Wray et al., 2008; Bishop et al., 2008a;

McKeown et al., this issue; Noe Dobrea, this issue] such as might be expected to be produced by a top-down leaching process. However, elsewhere, source regions for valley networks are not especially enhanced in phyllosilicate minerals [Bibring et al. 2006], and phyllosilicates associated with fluvial-lacustrine deposits discovered so far are most likely transported, not precipitated in-situ [e.g. Mustard et al., 2008; Grant et al., 2008;

Ehlmann et al., 2008b]. An alternative hypothesis to surface formation of phyllosilicates by weathering is that the formation of phyllosilicates occurred primarily in the subsurface, possibly driven by hydrothermally-mediated crustal cooling [e.g. Parmentier et al., 2008]. The intense meteoric bombardment of early Mars may also have played a role in phyllosilicate formation, perhaps in serving as a heat source generating liquid water from a cryosphere [e.g. Newsom, 1980]; impact craters on Earth commonly host hydrothermal deposits with phyllosilicates [Allen et al., 1982; Naumov, 2005; Nelson et al., 2005]. Key questions for addressing the timing, nature, and extent of aqueous alteration include: Which phyllosilicates are present? What other altered and unaltered minerals are present? What is the geomorphic setting and distribution of these diverse phases? And, how much phyllosilicate is present?

The question of abundance is difficult to address due to multiple scattering of photons in the visible and near-infrared wavelengths leading to nonlinearities in the relative contributions of components in mixtures, although initial efforts with OMEGA data have begun to address this question [Poulet et al., 2008]. This paper focuses on the methodology for identification of diverse hydrated silicates, along with associated silicate hydroxides and carbonates, and mapping of their occurrences using CRISM NIR data. 52

With the exception of the aluminum smectite montmorillonite, all non-sulfate hydrated minerals detected to date from orbit can be found in the region west of the Isidis basin, in and around the Nili Fossae. The remarkable mineralogic diversity of this region is probably due to a combination of factors: (1) a geologic history involving apparently multiple instances of alteration by liquid water, (2) the availability of starting materials of distinct mafic compositions, and (3) the action of erosion in producing excellent bedrock exposures.

After briefly reviewing the history and geologic setting of the study area west of

Isidis (Section 2.0) and methods of data analysis (Section 3.0), we detail in Section 4.0 the spectral data used as the basis for the detection of each hydrated mineral phase—Fe and Mg smectites, chlorite, prehnite, serpentine, kaolinite, K mica, the Na zeolite analcime, hydrated silica, and carbonate. We then describe in Section 5.0 the distribution and geomorphic setting of the mineral phases found in the region around the Nili Fossae.

By studying the geologic setting and associations of minerals, the nature of the formation environment can be partially constrained. We discuss the aqueous environments indicated, constraints on timing, and key outstanding questions from our study in Section

6.0. A companion paper provides an updated, comprehensive geologic history of the circum-Isidis region, examining the detailed stratigraphy of various mineral-bearing units, including mafic minerals [Mustard et al., this issue]. We focus here specifically on the identification of alteration minerals and on the various environments of aqueous alteration that can be ascertained, including evidence for low-temperature metamorphic or hydrothermal alteration in the subsurface.

2.0 Geologic setting, study area

West of the 1900-km Isidis basin, the cratered, Noachian-aged crust is cut by a series of concentric fractures formed by normal faulting in response to structural readjustment that shortly followed the Isidis impact around ~3.9 Ga (Figure 1, 2)

[Wichman and Schultz, 1989; Greeley and Guest, 1987; Werner, 2005]. These downdrop graben, collectively referred to as the Nili Fossae, extend from 20ºN to 26ºN, are up to 600 m deep and, along with the surrounding terrain, have been heavily modified by erosion and infilling. To the south, younger volcanic lavas from Syrtis Major overlie the

Noachian terrain and partially fill some of the fossae [Greeley and Guest, 1987]. The contact between the Hesperian Syrtis Major lavas and the Noachian etched plains marks a distinct compositional boundary between two crustal units (Figure 1). Generally the

Syrtis Major lavas are enriched in high-calcium pyroxene, while low-calcium pyroxene- bearing materials are more common in the Noachian cratered terrains [Mustard et al.,

2005]. The eastern portion of the Nili Fossae region also is the largest exposed olivine- bearing region on the planet [Hamilton et al., 2003; Hoefen et al., 2003; Hamilton and

Christensen, 2005; Mustard et al., 2005]. Although much of the olivine is in the form of sand, the olivine-bearing bedrock from which it is sourced is a contiguous unit that crops out over hundreds of square kilometers [Hamilton and Christensen, 2005; Mustard et al., this issue]. This bedrock unit likely originated from materials ejected during the Isidis impact [Mustard et al., 2007; this issue] or, alternatively, may represent low viscosity lava flows [Hamilton and Christensen, 2005; Tornabene et al., 2008].

Our study area encompasses the Noachian cratered terrain east from Nili Fossae westward to the eastern margin of the Antoniadi basin and from the northern Nilo-Syrtis 54 chaos southward to craters on the northern margin of the Hesperian Syrtis Major formation (Figure 2, 3). Both the Hesperian and Noachian units have sapping channels, fluvial valleys, and depositional fans that provide geomorphic evidence for sustained surface water flow [Mangold et al., 2007]. The largest such system is the watershed of

Jezero crater, which drained 15,000 km2 during the late Noachian or early Hesperian to form a paleolake with deltaic deposits [Fassett and Head, 2005; Ehlmann et al., 2008b].

However, while the Noachian terrain shows extensive mineralogic evidence of aqueous alteration, discussed further in this paper, the Hesperian Syrtis Major lavas do not show evidence in orbital data of hydrated minerals [Mangold et al., 2007].

3.0 Datasets and Methods

3.1 Data processing and selection of spectra

CRISM is a hyperspectral imaging spectrometer on the Mars Reconnaissance

Orbiter (MRO) with 544 channels that sample the visible to near-infrared spectral region from 0.4 - 4.0 μm using two detectors [Murchie et al., 2007a]. The S detector samples wavelengths 0.4-1.0 μm, and the L detector samples wavelengths 1.0-4.0 μm. CRISM operates in two modes: (1) a 72-channel mapping mode that will provide global coverage at 200 m/pixel and (2) a full 544-channel targeted mode that provides 10 x 10-20 km images at a resolution of 15-38 m/pixel. Results from the L detector data of the latter, high-resolution mode are the focus of this paper and allow mapping of mafic minerals and hydrated or hydroxylated alteration minerals.

Spacecraft data were converted to I/F as described by Murchie et al. [2007a; this issue a]. Prior to spectral analysis, simple atmospheric and photometric corrections were 55 applied to CRISM data cubes to correct for viewing geometry and separate out the contribution of reflected light from the surface from that of the atmosphere. Cubes were photometrically corrected by dividing by the cosine of the incidence angle, assuming that to first order the surface behaves as a Lambertian scatterer. The contribution of the atmosphere was removed using the volcano scan correction method, also successfully employed by the OMEGA mission [Mustard et al., 2005]. In the volcano scan method, an atmospheric transmission spectrum is derived from observations at the base and top of

Olympus Mons. For a given CRISM observation the atmospheric transmission spectrum is scaled to match the band depth of the 2-μm CO2 feature in each pixel, and then the observed spectrum is divided by this scaled spectrum. The technique, which assumes a multiplicative atmospheric contribution to surface spectra, removes absorption features resulting from atmospheric gases with few residual artifacts, but aerosols and ices are not accounted for and may lead to spectral slopes or residual absorptions. Overall, the volcano scan technique is simple and efficient for correcting large numbers of CRISM scenes. Results compare favorably with the more rigorous, but time-intensive scene- specific radiative transfer modeling employed using DISORT [Arvidson et al., 2006] provided that aerosol abundances at the time of imaging are not appreciably different from those present at the time of the volcano scan observation. Nevertheless, small positive features or residual absorptions near 2 μm are still commonly present even in atmospherically corrected CRISM data (as is apparent in some CRISM spectra from

Figures 4-12). These are artifacts of the atmospheric calibration methods rather than features of use in mineral identification. 56

Following atmospheric correction, a noise removal algorithm that removes data spikes in both the spectral and spatial domains but does not affect broader absorptions of mineralogic interest was implemented [Parente, 2008]. In addition, to highlight spectral differences between areas and reduce residual atmospheric and instrumental artifacts in spectra, average spectra from regions of interest were divided by spectra from a dusty or spectrally unremarkable region in the same scene. Where possible, n x n pixel spectra of similar albedo to the n x n pixel spectrum of interest and from the same column of the unprojected image were used as the denominator. Sampling in the same column reduces detector-dependent noise. Alternatively, larger, more irregularly shaped regions of interest for the numerator and denominator were identified, not necessarily from the same column. In both cases, tens to thousands of pixels were averaged to minimize spurious absorptions due to noise (Appendix Table 1). Even with these steps, some artifacts persist. After removing CRISM known bad spectral channels, some known residual artifacts are common near 2.0 μm due to errors in atmospheric calibration and at 1.65 μm due to a detector filter boundary [Murchie et al., this issue a]. Single channel spikes in spectra that are too narrow for absorptions related to mineralogy are also removed or ignored. Spectra used in the detailed analysis of Section 4.0 are either from the central portion of the scene or have been corrected for column dependent wavelength shifts due to the CRISM spectral smile by using the CDR WA file [Murchie et al., 2007a; Murchie et al., 2007b]. Band centers reported here from the CRISM data have an error of approximately ±0.01 μm, the average spectral sampling interval of the instrument.

3.2 Mineral identification and mapping 57

Hydrated silicate minerals are identified in CRISM spectroscopic data by overtones and combinations of fundamental vibrational absorption features in the region from 1.0-2.6 μm. The presence of H2O in the mineral structure is expressed near 1.9 μm by a combination tone of the fundamental bending and stretching vibrations of the water molecule. Structural OH stretching vibrations occur near 1.4 μm, as do overtones of the

H2O molecule and combinations of one stretching plus two bending vibrations of the H2O molecule [Bishop et al., 1994]. The position of the OH overtone shifts depending on the type of octahedral cation (Mg, Al, Fe) from 1.38 to 1.43 μm for many phyllosilicates [e.g.

Bishop et al., 2002a; 2002b]. From 2.1-2.5 μm, additional structural OH combination stretching plus bending vibrations are found, and their precise wavelength position depends on the type of octahedral cation and mineral structure. For example, Fe-OH absorptions in smectites are found near 2.28 μm whereas Al-OH absorptions are found near 2.20 μm [Clark et al., 1990]. Overtones/combinations of fundamental vibrations related to H2O in zeolites [Cloutis et al., 2002] and CO3 in carbonates [Gaffey, 1987] are also found in the 2-2.6 μm wavelength region.

These absorptions can be parameterized and mapped spatially, e.g. by calculating absorption band depth for each pixel in an image. Here, we utilize the standard parameter set of Pelkey et al. [2007] formulated for multispectral CRISM images along with custom parameters for absorptions at 1.9, 2.2, and 2.5 m. A 1.9-m absorption occurs in minerals with H2O, a 2.2-m absorption occurs in minerals such as Al phyllosilicates and opaline silica with Al-OH or Si-OH, and a 2.5-m absorption occurs minerals such as in carbonates and zeolites. The devised custom parameters take advantage of the additional bands available in hyperspectral targeted images to average out noise and reduce spurious 58 high values or, in the case of BD2500, provide additional parameters to the standard multispectral CRISM set and are formulated as follows:

BD1900H = 1 – AVG(R1908:R1914)/(AVG(R1862:R1875)+AVG(R2112:R2126))

BD2200 = 1 – 2*AVG(R2199, R2205)/(AVG(R2132, R2146)+AVG(R2252+R2258))

BD2500 = 1 – AVG(R2500, R2510)/AVG(R2380, R2540) where, for example, R1980 is the reflectance at 1.980 m, and avg(R1908:R1914) is the average reflectance over all wavelengths between 1.908 and 1.914 m. The other commonly employed parameters from the standard Pelkey et al. [2007] set were D2300, used for mapping absorption bands near 2.3 m (common in Fe/Mg smectites and carbonates), and BD2350 used for mapping specifically the 2.35-m absorption band

(common in chlorites, prehnite, illite/muscovite, and some carbonates). Generally, threshold values used in mapping are 0 to 0.05. However, factors other than absorption band strength from minerals of interest can affect parameter values computed, e.g. detector noise, aerosols, continuum slope; so threshold values were manually adjusted on an image-by-image basis by analyzing the spectral data to verify that parameter maps were indeed mapping out occurrences of spectra corresponding to definitive mineral identifications. Actual lower bounds used ranged from -0.005 to 0.01 and upper bounds from 0.02 to 0.1. Computed CRISM image data and parameter maps were then map- projected using the line-of-sight intercept of each detector element with the 128-pixel per degree MOLA shape model of Mars contained in band 10 of the Derived Data Record

(DDR) accompanying each image [Murchie et al., 2007b; this issue a].

In order to understand the geomorphology and geologic setting of the mineral detections, these CRISM parameter maps were coregistered with and combined with a 59 variety of imaging datasets. These include coordinated observations acquired by the two high-resolution MRO cameras, the Context Imager (CTX) and the High Resolution

Science Experiment (HiRISE) cameras on MRO, which acquire images at 5 m/pixel and

0.25 m/pixel, respectively [Malin et al., 2007; McEwen et al., 2007]. Additionally,

Thermal Emission Imaging System (THEMIS) day infrared images at 100 m/pixel were sometimes employed [Christensen et al., 2004]. Mars Orbiter Laser Altimeter elevation data were used to view three dimensional representations of the topography. These data were obtained for individual images from their DDRs and from the 128 pixel per degree global map for regional topographic information.

Full- and half-resolution CRISM images were targeted to follow up on detections of altered and unaltered minerals using coarser spatial resolution OMEGA and CRISM mapping data, to focus on areas where prior datasets indicated exposed stratrigraphy or interesting morphologic features, and as ride-alongs to HiRISE and CTX imaging requests. Placement of CRISM targeted images reflects an attempt to ascertain geologic relationships between mineralogic units exposed from beneath dust and capping materials with no distinctive NIR spectral signature rather than a systematic survey of the study area.

4.0 Identified minerals

Most of the numerous diverse hydrated silicate minerals recently identified on

Mars by CRISM can be found in the region around Nili Fossae [Mustard et al., 2008].

Based on CRISM VNIR spectra, below we detail the rationale for the identification of multiple varieties of Fe/Mg smectite, chlorite, prehnite, serpentine, kaolinite, a K mica 60 such as illite or muscovite, hydrated silica, the zeolite analcime, and magnesium carbonate. We also describe spectra consistent with either a mixture of silica and Mg sulfate or silica and zeolite. Locations of CRISM observations from which spectra were derived are shown in Figure 3. Tables in the Appendix provide the locations and number of pixels averaged for all CRISM spectra as well as complete identifying information for library spectra used in Figures 4-12. The intent of Section 4.0 is to detail the basis for these phases’ unique identification so that they may be replicated by future users of

CRISM data and also to provide detailed discussions of mineral phases and uncertainties of interest for geochemical modeling. The discussion of the geologic context of these minerals begins in section 5.0. Three guiding principles used in our mineral identifications were that (1) absorption features in the CRISM spectra should be consistent with the same features in laboratory spectra of the mineral, (2) these features should not be consistent with those of other minerals (for uniqueness), and (3) no strong absorptions should be left unexplained.

4.1 Fe/Mg smectites

As previously revealed by the OMEGA spectrometer, iron magnesium smectites are the areally dominant phyllosilicate in the Nili Fossae region [Poulet et al., 2005;

Mangold et al., 2007]. Fe/Mg smectites are 2:1 phyllosilicates and form from weathering and hydrothermal processes. They are identified in spectroscopic data by a characteristic vibrational absorption near 2.3 μm and a 1.9-μm band indicating molecular H2O (Figure

4). A 1.4-μm band is commonly discernible as are additional combination tones near 2.4

μm. CRISM targeted observations reveal a suite of spectra with these characteristic features in the Nili Fossae region but which show considerable variation both within and 61

between scenes. A subset of the diverse spectra we classify as Fe/Mg smectite is shown

in Figure 4. We focus here on determining with greater precision the likely chemistry of

the various Fe/Mg smectites in Nili Fossae.

A variety of smectite compositions are possible depending on Al, Fe, and Mg

occupancy of the octahedral sheet and substitution of cations in the interlayer to balance

+ 2+ 3+ the charge. Smectites have a general formula of M x-y+2z(R 3-y-z, R y, z) Si4- xAlxO10(OH)2 where 0.3 x-y+2z 0.6 and z refers to the number of vacancies per unit cell in the octahedral layer [Meunier, 2005]. There is continuous solution between Fe and

Mg smectite endmembers [Grauby et al., 1994], and compositional variation is expressed in variations in the NIR spectra. In particular, the relative proportions of Fe, Mg, and Al cations in the octahedral site dictate the presence and position of bands from 1.35-1.45

μm and from 2.0-2.5 μm. That the smectites in Nili Fossae are predominantly Fe/Mg rather than Al is determined by the presence of an absorption near 2.3 μm rather than near 2.2 μm. The 2.3-μm band results from combination tones of the 2Fe-OH and 3Mg-

OH bends and stretches, and a more subtle band at 2.4 μm is also related to these bonds

[Clark et al., 1990; Frost et al., 2002; Bishop et al., 2002b]. The precise position of the

2.3-μm band varies depending on the relative proportions of Fe vs. Mg. For example, dioctahedral nontronite, the Fe endmember, has a band center near 2.29 μm [Grauby et al., 1994; Bishop et al., 2002b; Frost et al., 2002], whereas the trioctahedral Mg endmembers saponite and hectorite have absorptions near 2.31-2.32 μm (Figure 4b)

[Clark et al., 1990]. Aluminum, if present in smectite, is indicated by a strong 2.21 μm

2Al-OH combination tone or a somewhat weaker Al,Fe-OH combination tone at 2.23-

2.25 μm [Bishop et al., 2002a, 2008]. 62

In the Nili Fossae region, the 2.3-μm band varies in shape and strength from a sharp absorption of up to 10% band depth to a broad L-shaped shoulder (e.g. 50F2 c/m and A053 r/g in Figure 4). Most Fe/Mg smectites in the Nili Fossae region exhibit an absorption centered at 2.30 μm, intermediate between the Fe and Mg endmembers (e.g.,

3E12 r/g and 9D44 c/m, Figure 4). This contrasts with the mostly nontronite-like spectra that are areally dominant in and around Mawrth Vallis [Noe Dobrea et al., this issue]. Yet spectra that exhibit absorptions at 2.29 μm are not uncommon, found mostly in craters to the west of Nili Fossae and in Nilo Syrtis (e.g. 9312 c/m, 527D respectively). In many cases, these are paired with weaker absorptions near 2.4 μm. Spectra with absorptions near 2.32 μm are also found but less commonly (e.g., 64D9 r/g).

A defining feature of smectites is that molecular water in the interlayer is a necessary part of the mineral structure. The strength of the ~1.9-μm band in spectral data is related to the degree of hydration, both in water molecules in the interlayer and surface hydration; this can vary depending on the thermal history of the smectite and the nature of the water [Milliken and Mustard, 2005]. Dehydration may be the cause of reduction in the strength of the ~1.9 band relative to the ~2.3-μm band in data from Nili Fossae.

Additional intermixed hydrated phases, e.g. zeolite, may also enhance the 1.9-μm band without introducing additional bands in the wavelength region from 2.0-2.6 μm.

Although the strength of the 1.9-μm band in Nili Fossae spectra shows considerable variation, the position of this band does not vary among the Fe/Mg smectites either in

Nili Fossae or in laboratory data so it does not provide compositional information.

On the other hand, there is significant variation in the strength and position of the

1.4-μm band within Fe/Mg smectites both at Nili Fossae and in laboratory data. 63

Combination tones of H2O are present at 1.41 μm as well as the structural OH stretching overtone that depends on the octahedral cation [Bishop et al., 1994]. This stretching overtone occurs at 1.43 μm for 2Fe-OH [Bishop et al., 2002a,b; Frost et al., 2002], at

1.41 for 2Al-OH [Bishop et al., 2002a,b], and at 1.38-1.39 μm for 3Mg-OH [Clark et al.,

1990; Bishop et al., 2002a, 2002b]. In the Nili Fossae region, the ~1.4-μm band varies in position from 1.39 to 1.43 μm, indicating a range of phyllosilicate chemistry. Notably, in the more nontronite-like spectra with absorptions at 2.29 μm, the OH overtone band is centered near 1.43 μm, as would be expected to result from a mostly iron-bearing smectite with dominantly Fe2-OH bonds in the octahedral sites. In the most common Nili

Fossae smectites with band centers from 2.30-2.31 μm, the greater proportion of Mg is indicated by the shift shortward of the ~1.4-μm band. In some spectra, the 1.4-μm band is weak to absent, which may be due to the effect of mixing with and/or coating by more opaque iron-bearing oxides, which commonly occur in association with Fe/Mg smectite clay minerals and can obscure the ~1.4-μm bands [Swayze et al., 2002; 2003].

The data collectively suggest diverse Fe/Mg smecites in the Nili Fossae region.

Dominantly Mg smectites such as the trioctahedral smectites saponite or hectorite are indicated in most locations. Another Mg phyllosilicate, the fibrous clay sepiolite, cannot be excluded but rarely forms on Earth in large 100s of km-scale deposits such as observed in Nili Fossae. Additionally, more Fe-rich smectite clays are common in some parts of the Nili Fossae region, and these are likely nontronites. Notably, there is no spectral evidence for aluminum comprising an appreciable proportion of cation sites in

Nili Fossae Fe/Mg smectites. While the relatively weak Fe2OH and Mg3OH bands near

2.4 μm can be discerned, neither a 2.21 μm (2Al-OH) or 2.23-2.25 μm (Al-Fe-OH) 64 absorption related to Al has been identified as would be found in montmorillonite or in ferruginous smectites with significant aluminum (e.g. SWa-1, Figure 4b).

4.2 Chlorite and Prehnite

In nearly a dozen CRISM images, mostly associated with craters west and south of Nili Fossae, a spectral class is detected with a prominent vibrational absorption centered at 2.34-2.35 μm with a shoulder at 2.26 μm. The 1.9-μm band is usually weaker than in smectites or absent, and there is typically a weak but sharp 1.40-μm band and, in some cases, an additional sharp band near 1.48 μm (Figure 5). We identify two minerals as being responsible for spectra of this character: an Fe-rich chlorite and prehnite.

Chlorites are 2:1:1 phyllosilicates with a T-O-T layer structure and a brucite-like octahedral interlayer and can form from primary (magmatic), hydrothermal, metamorphic, and diagenetic reactions. Bands from 2.2-2.5 μm vary depending on substitution of cations such as Al, Fe, and Mg in octahedral and tetrahedral layers. The strong absorption at 2.33-2.35 μm results from a combination of overtones of the Mg-OH stretch and bend and the 2.25-2.26 shoulder results from Al,Fe-OH or Al,Mg-OH stretching modes. This latter band, along with a greater band width at 2.33-2.35 μm distinguishes chlorite from serpentine [Bishop et al., 2008]. In chlorite, the band at 2.33-

2.35 μm shifts to longer wavelengths as the proportion of iron relative to magnesium in the chlorite increases (Figure 5c) as has been reported for other chlorite bands [King and

Clark, 1989]. In the Nili Fossae region, CRISM spectra have a strong 2.35-μm absorption, indicating the chlorite is comparatively iron-rich (Fe:Mg>~0.5). A ferroan

2+ clinochlore (e.g. ripidolite; (Mg, Fe )5Al(Si3Al)O10(OH)8) [Post and Plummer, 1972] provides the best match to the CRISM spectra (e.g., for 50F2 s1b/s2b in Figure 5). 65

Prehnite is a calcium aluminum silicate hydroxide and forms from hydrothermal or metamorphic activity under specific conditions of <3 kbar, 200-350°C, and XCO2 <

0.004. It typically forms in association with chlorite and pumpellyite [Frey and Robinson,

1999]. Previously, Clark et al. (1990) identified the spectral features of prehnite in ground based spectroscopic observations of Mars as those of scapolite but recently recognized that absorptions at 1.48 and 2.36 μm were due to prehnite contamination in their laboratory scapolite samples [Clark et al., 2000; Clark et al., 2008]. Prehnite occurs in multiple locations in ancient Martian cratered terrains [Clark et al., in prep], including the greater Argyre basin region [Buczkowski et al. this issue]. The band position of 2.35-2.36

μm measured in laboratory spectral data of prehnite coincides with that of Fe-rich chlorite (Figure 5c). However, the presence of an absorption at 1.48 μm appears, from examination of numerous spectral libraries, to be uniquely diagnostic of prehnite and is the basis for identification of this mineral in the Nili Fossae region. Some CRISM spectra have absorptions at 2.35 μm, weak to absent 1.9-μm bands, and absorptions at both 1.41 and 1.48 μm (Figure 5). We infer this spectrum to result from a mixture of phases bearing prehnite. A likely additional component in that mixture is chlorite. Pumpellyite also has a similar spectral character to chlorite, in particular more Mg-rich chlorite, but to date we do not observe a band center at 2.33 μm in the Nili Fossae spectra, so this phase cannot be positively and uniquely identified.

4.3 Serpentine

A spectral class that occurs infrequently in the Nili Fossae region (two CRISM images to date) is distinguished by an absorption centered at 2.32 μm with additional 66 absorptions at 2.11 μm and 2.51 μm and a sharp vibrational feature at 1.39 μm (Figure 6).

We identify this mineral as a magnesium-rich serpentine mineral.

Serpentines are 1:1 trioctahedral phyllosilicates with the general formula (Fe,

Mg)6Si4O10(OH)8 and form during metamorphism, hydrothermal activity, or weathering of ultramafic rocks. They have numerous combination tones of metal-OH bends and stretches from 2.3-2.7 μm [King and Clark, 1989; Post and Borer, 2000]. The strongest of these is centered at 2.32 μm for Mg-OH in magnesium serpentines (Figure 6c). This band has a distinctly different appearance in Fe-rich serpentines, where it occurs at longer wavelengths and is in some cases accompanied by aluminum bands at shorter wavelengths [Calvin and King, 1997]. The band at 2.50-2.51 μm is the next strongest of these features. Mg serpentines also have a diagnostic band at 2.10-2.12 μm. Serpentines have an OH stretching overtone located at 1.39 μm for Mg serpentines that shifts to somewhat longer wavelengths (1.40-1.41 μm) for Fe-rich serpentines [King and Clark,

1989; Calvin and King, 1997].

In the CRISM spectra, from 2.0-2.6 μm, we observe a sharp absorption at 2.32

μm and additional absorptions at 2.51 and 2.11 μm (Figure 6). There is a 1.9-μm band indicating the presence of H2O that is broad in the 2-μm region. There is an additional sharp vibrational band at 1.39 μm. All of these features are consistent with the presence of an Mg serpentine such as antigorite, lizardite, or chrysotile. These can be distinguished from the Mg smectite, saponite, which also has a band at 2.32 μm, by the sharpness and position of the 1.39-μm band and the additional band at 2.10 μm. Stoichiometric serpentine does not have molecular water in its structure, which would create the 1.9-μm 67 band so we are likely observing Mg serpentine in a mixture with a hydrated phase in Nili

Fossae.

4.4 Kaolinite

In over a dozen CRISM images in and around Nili Fossae, a spectral class is detected that has prominent asymmetric absorptions at 1.4 and 2.2 μm (Figure 7). The

2.2-μm feature commonly occurs as a doublet with a second absorption near 2.16 μm. A

1.9-μm absorption is present but shows significant variation in strength. We identify this spectral class as a kaolinite group mineral. The doublet at 2.16 and 2.21 μm is due to a combination of vibrations of 2Al-OH groups in the mineral structure [Clark et al., 1990].

The presence of a doublet at 2.2 allows kaolinite to be distinguished from the Al-smectite montmorillonite where only a single absorption is present at 2.2 μm [Petit et al., 1999].

The sharp absorption near 1.4 μm with its distinctive shoulder is also diagnostic. The particularly sharp 1.41-μm absorption in kaolin group minerals is caused by vibrations of inner hydroxyl groups between the tetrahedral and octahedral sheets leading to multiple minima near 1.4—a triplet in well-crystalline samples [Crowley and Vergo, 1988; Clark et al., 1990; Bishop et al., 2002].

Kaolinite group minerals are 1:1 phyllosilicates with the formula Al2Si2O5(OH)4 and include kaolinite and the rarer dickite, nacrite and halloysite, the latter of which is a hydrated form [Giese, 1988]. Some features in IR spectra can be used to differentiate between these similar minerals. In particular, dickite and nacrite have a prominent short- wavelength absorption in the doublet at 2.17-2.18 μm whereas kaolinite and halloysite have a weaker absorption at 2.16 μm as is observed in the CRISM data from Nili Fossae

(Figure 7). Kaolinite forms as a weathering product, via hydrothermal alteration, or as an 68 authigenic sedimentary mineral [Murray, 1988]. The hydrated form, halloysite, usually forms under weathering or hydrothermal processes that favor a more disordered structure, such as alternating wet and dry conditions or rapid growth [Giese, 1988]. Well-crystalline kaolinite has a very sharp, defined doublet with two distinct absorptions at CRISM spectral resolution. Disordered kaolinite, halloysite, and kaolinite-smectite mixed layer phyllosilicates have less well-defined doublets as can be seen in the lower panel of Figure

7c [Crowley and Vergo, 1988]. CRISM spectra are generally more similar to these forms of less well-crystalline kaolinite and have a 1.9-μm band indicating structural water as is typical for halloysite and kaolinite-smectite mixed phyllosilicates. However, the separation of the doublet absorptions is near the level of noise in the CRISM data and spatial averaging to reduce noise tends to decrease the prominence of the doublet. It is also likely that a pixel might include other hydrated minerals, spatially mixed at a subpixel scale, which might contribute to the 1.9-μm band. The detection of a kaolinite group mineral is definitive, and a poorly crystalline kaolinite, halloysite or a mixed layer kaolinite-smectite (with kaolinite > smectite) are the likely specific mineral candidates.

4.5 Illite/Muscovite

In five CRISM images in the vicinity of a 50 km impact crater west of Nili

Fossae, a spectral class is detected with sharp vibrational absorptions at 1.4, 2.2 and 2.35

μm. The 1.9-μm absorption is weak, and in some cases a weak 2.44-μm absorption is present (Figure 8). In these scenes, when the 2.2-μm feature is present, it is always accompanied by a 2.35-μm band. The strength of the 2.2-μm feature is directly proportional to the strength of the 2.35-μm band, making it unlikely the phase is a mixture. 69

We identify this phase as a probable potassium mica. The presence of the bands at

2.2 μm, 2.35 μm and 2.44 μm distinguish K micas, i.e. illite and muscovite, from montmorillonite, which has a single absorption at 2.2 μm [Clark et al., 1990].

Additionally, the sharpness of the 1.4-μm absorption and the weak to absent 1.9-μm band indicate a phyllosilicate with little to no structural H2O. Illite and muscovite are high layer charge (0.6-1.0), non-expanding 2:1 phyllosilicates that contain primarily Al in the octahedral layer, some Al in the tetrahedral layer, K in the interlayer region [Meunier,

2005]. The absorption features from 2.2-2.5 μm in K micas are due to stretch plus bend combination bands of structural OH bound to Al [Clark et al., 1990].

Distinguishing between illite and muscovite is challenging. Illite is a K-deficient intermediate phase between the micas pyrophyllite (Al2Si4O10(OH)2) and muscovite

(KAl2AlSi3O10(OH)2). Muscovite can be magmatic, hydrothermal, or metamorphic. Illite can form from hydrothermal, diagenetic, or weathering reactions. In diagenetic or weathering reactions, disordered or 1M polytypes of illite with significant interlayer water or mixed layer illite-smectite clays are the most common [Rosenberg, 2005]. With increasing temperature, pressure, or time, illite undergoes a polytypic conversion of 1Md to 1M to 2M1 and eventually to muscovite [Lonker and Gerald, 1990; Rosenberg, 2002].

NIR vibrational absorptions in spectra of these various polytypes of illite and muscovite usually differ in their relative strengths, though not positions (Figure 8c) [Clark et al.,

1990; Kruse and Hauff, 1991].

Because of the small size (<15 CRISM pixels at most) of the K mica-bearing exposures found to date, obtaining “pure” spectra of this phase is difficult. Chlorite and/or prehnite are found in the same images and in geomorphically similar knobby, 70 ridged terrain. Their absorptions at 2.35 μm obviously contribute to the strength of the

2.35-μm feature in some K mica-bearing spectra, and the 1.48-μm absorption probably indicates the presence of prhenite (e.g. 454E in Figure 8). The 2x2 average spectrum from HRS00002FC5 is the “purest” found to date. The sharpness of the 1.4 and 2.2-μm absorptions, the lack of a 1.9-μm absorption, and the distinguishable 2.35 and 2.44-μm absorptions indicate K mica. The small size of exposures imaged to date and consequently relatively high noise in spectra along with the similarity of illite and muscovite in VNIR spectral data preclude a more specific identification.

4.6 Zeolite: analcime

In several CRISM images west of Nili Fossae in craters near the Antoniadi basin, spectra from small (tens of pixel at most) regions of interest have strong absorptions at

1.42 μm, 1.91 μm, and 2.52 μm and a weak absorption at 1.78 μm (Figure 9). We identify this phase as the zeolite mineral analcime.

Zeolites are aluminum silicates that are made up of chains of rings of SiO4 and

AlO4 tetrahedra. Water molecules and other cations occupy the centers of cages in the framework structure. Analcime in particular is a sodium-rich zeolite with a formula

NaAlSi2O6•H2O. It forms under hydrothermal conditions, from weathering and diagenetic reactions of igneous minerals, particularly volcanic glass, in alkaline, saline waters, and has been hypothesized to infrequently occur as a primary magmatic product [Hay, 1986;

Line et al., 1995; Luhr and Kyser, 1989]. NIR spectral features of zeolites are dominated by sorbed water [Hunt and Salisbury, 1970; Clark et al., 1990]. However, laboratory spectra of analcime are quite distinct from those of other zeolites, even Na-rich ones such as natrolite [Clark et al., 1990; Cloutis et al., 2002] (Figure 9c). Analcime is characterized 71 by strong absorptions at 1.42 μm, 1.91 μm, 2.52 μm and progressively weaker absorptions at 1.79 μm, 1.12 μm, 2.14 μm. The 1.79 μm band is seen in CRISM data, but the latter two are not. This may simply be a function of abundance—the weakest bands are the first to “disappear” at the SNR level as abundance decreases—or a result of Fe- related minerals obscuring the shorter wavlength features (as discussed in Section 4.1 for smectites). The spectral shape and absorption band positions in the CRISM data closely match those of analcime, whose suite of absorptions in this wavelength range are apparently unique among minerals measured by NIR spectroscopy to date.

4.7 Hydrated silica

In over a dozen CRISM images taken within craters in the Nili Fossae region and over topographic ridges within the Antoniadi basin, a phase occurs with a broad absorption at 2.2 μm and also absorptions at 1.4 μm and 1.9 μm (Figure 10). We identify this phase as hydrated silica, which has been previously described in detail for layered deposits near Valles Marineris [Milliken et al., 2008].

Opaline silica and hydrated (H2O-bearing), hydroxylated (OH-bearing) basaltic glass exhibit absorption bands due to combination tones of Si-OH at 2.21-2.22 μm due to isolated Si-OH and at 2.26 μm due to H-bound Si-OH [Anderson and Wickershiem,

1964; Stolper, 1982; Goryniuk et al., 2004]. These bands near 2.2 μm are clearly wider than those observed in aluminum phyllosilicates [Milliken et al., 2008]. The bands at 1.91

μm are due to H2O in the mineral structure whereas those near 1.4 μm result from both

H2O and structural OH. The 1.4-μm band position in both opal and hydrated basaltic glass is near 1.40-1.41 μm [Langer and Florke, 1974; Stolper, 1982], although this can shift shortward to as much as 1.38 μm when only the influence of the Si-OH bond 72

remains, following the removal of H2O [Anderson and Wickershiem, 1964; Swayze et al.,

2007; Milliken et al., 2008]. With less water, the influence of the 2.26-μm band is also diminished resulting in a slight narrowing in the 2.2-μm band.

In Nili Fossae, we observe spectra with a broad absorption centered near 2.21 μm with a 1.41-μm band position as well as spectra with a broad but slightly narrower absorption at 2.21 μm with a weaker 1.9-μm band, and a 1.39-μm band (Figure 10c).

These spectra are consistent with hydrated silica with varying amounts of H2O in the structure. The former is more consistent with opal—hydrated amorphous silica; the latter is more consistent with a partially dehydrated opal or altered basaltic glass.

4.8 Other zeolites or polyhydrated sulfate mixtures

In four CRISM images to date from the Nili Fossae region, spectra have been collected with stronger water-related bands at 1.43 μm and 1.92 μm than commonly observed in most other Martian terrains bearing hydrated minerals (Figure 11). These spectra also display very subtle absorption features at 2.20 μm, 2.26 μm, and 2.42 μm.

As discussed above, absorptions near 2.20 μm and 2.26 μm may indicate the presence of opaline phases or hydrated, hydroxylated glasses. However, the band positions in the CRISM data are at 1.43 μm and 1.93 μm, i.e. shifted longwards from typical locations for these bands in hydrated silica. Along with the weakness of the bands near 2.2 μm, this suggests a mixture of phases for this spectral class. The dominance of water-related bands in the spectrum and the absence of other diagnostic bands precludes a precise mineral identification at present. As discussed above, NIR spectral features of zeolites are dominated by sorbed water [Hunt and Salisbury, 1970; Clark et al., 1990] as are spectral features in polyhydrated magnesium and ferric sulfates [Cloutis et al., 2006] 73

(Figure 11c). The positions of the 1.43-μm and 1.92-μm bands in the CRISM data best match those of a zeolite. Absorptions near 1.4 and 1.9 μm in sulfates are usually at longer wavelengths. However, in a sulfate-opaline silica mixture, broader bands with centers intermediate between sulfate and silica might be expected. A very subtle inflection is seen in CRISM data at 2.4 μm. This absorption occurs in the Group 5 structural class of zeolites and may be an H2O combination band [Cloutis et al., 2002] but also occurs in spectra of some polyhydrated Fe and Mg sulfates [Cloutis et al., 2006]. While the mineral identification for this spectral class of hydrated material is currently indeterminate, the most likely candidates are either a silica-zeolite or a silica-sulfate mixture.

4.9 Carbonate

In over two dozen images in the eastern portion of Nili Fossae, spectra with paired absorptions at 2.31 and 2.51 m are observed. These characteristic spectra also have a 1.9-m band and a broad absorption centered near 1.0 m (Figure 12). This phase is identified as magnesium carbonate as detailed by Ehlmann et al. [2008a].

Paired absorptions of equivalent strength at 2.3 and 2.5 m are overtones of the fundamental absorptions of carbonate that occur in the 6-15 m range [Lane and

Christensen, 1997], and the precise wavelengths of the minima depend on the cation paired with the carbonate [Hunt and Salisbury, 1971; Gaffey, 1987]. Magnesite or a mixture of hydromagnesite+magnesite provide the best spectral match to the band shape and position of the 2.31 and 2.51-m bands observed in CRISM spectra (Figure 12c).

The carbonate phase we see is in some cases clearly in a mixture and contains a hydrous component, either hydrous carbonate or another constituent. Iron magnesium smectites have a strong absorption at 2.3 m and successively weaker absorptions at 2.4 and 2.5 74

m (nontronite, Figure 12c; Figure 4). The presence of a small additional feature at 2.4

m, sometimes present with the carbonate (e.g. A09C, Figure 12b), probably indicates the presence also of Fe/Mg smectite, either spatially or intimately mixed. The broad absorption at 1.0 m is due to ferrous iron either in small amounts in the magnesite or in associated olivine [Ehlmann et al., 2008a].

In contrast to most hydrated silicates, carbonate also has diagnostic absorptions in the 3-4 m range, specifically broad absorptions at 3.4 and 3.9 m. CRISM data from

Nili Fossae in carbonate-bearing terrains show a strong 3.9-m absorption band and a considerably weaker 3.4-m absorption (Figure 12e) [Ehlmann et al., 2008a]. The 3.4 m band may be weakened or distorted by what is apparently an instrumental artifact centered at 3.18 m (Muchie et al., this issue a), which does not map coherently with geomorphic units either here or elsewhere on the planet and does not correspond to known absorptions in library spectra. The 3.4 and 3.9-m bands in carbonate are in some cases quite strong and of equivalent strength (e.g. Mag. CB-EAC-006A, Figure 12f); however, sometimes the 3.4 m band may be considerably weaker than the 3.9-m band

(e.g. Mag CC-JFM-006B and mag.+hydromag., Figure 12f). The natural variability of these band strengths is not well-studied. Grain size may play a role, and water can also act to subdue bands in the 3-4 μm range. In hydrous carbonates, bands from 3-4 m may be entirely absent (Figure 12f) [Calvin et al., 1994].

Unratioed CRISM spectra have a strong positive slope at wavelengths greater than ~3.5 m due to the thermal emission contribution of the surface. The 3.9-m band can be distinguished as an inflection, or slightly reduced steepness over wavelengths >3.7

m, in carbonate-bearing materials relative to non-carbonate-bearing materials (Figure 75

12d). The strongest 3.9-m absorptions map spatially in the same locations as the 2.31 and 2.51 m bands [Ehlmann et al., 2008a, Supplementary Online Material]. The 2.51-

m band never occurs alone in spectrally sampled terrains in this part of the planet but is always accompanied by the 2.31 m band. These spectral characteristics allow magnesium carbonate to be identified and mapped.

5.0 Geomorphology and Geologic Setting

5.1 Mineral distribution: provinces of alteration

With the higher spatial resolution provided by CRISM, hydrated minerals in the greater Nili Fossae region are detected extending further westward than previously known from analysis of OMEGA data [Mangold et al., 2007; Mustard et al., 2007].

Examination of Figure 2 shows that Fe/Mg smectites are the most common alteration mineral in the Nili Fossae region, as is true of Mars globally. Fe/Mg smectites are nearly ubiquitous in Noachian terrains. They are detected in targeted images over the entire study area except in the Hesperian Syrtis Major lava flows, dust-covered terrains north of

Baldet crater, and east of the Nili Fossae where the uppermost unit is Amazonian Isidis basin fill.

In contrast to the consistent presence of the Fe/Mg smectites, other hydrated minerals show considerable heterogeneity in distribution. Kaolinite and carbonate are found only in the area immediately around the Nili Fossae, with carbonate specifically restricted to the easternmost part of the area (Figure 2). In and around impact craters, especially to the west and south of Nili Fossae, different alteration minerals are observed, especially chlorite and prehnite. From crater to crater, there are distinctively different 76 mineralogic assemblages, i.e. combinations of mineral-bearing units detected in the same geographic area. For example, a 50 km crater (20°N, 69°E) has a prehnite-chlorite-K mica assemblage whereas two 25 km impact craters (19°N, 65°E) have a smectite- chlorite-hydrated silica-analcime assemblage.

Below, we consider in greater detail the geomorphology of the mineral-bearing units for each alteration mineral identified. In accordance with the distinctive mineralogic assemblages identified and their distinctive geomorphic settings discussed below, we divide the region into provinces of alteration: Eastern, Western, and Central/North Syrtis.

5.2 Eastern province: Nili Fossae stratigraphy

Like OMEGA data, CRISM data show Fe/Mg smectite-bearing rock units are the lowermost exposed stratigraphic unit, i.e. the basement rock, in eastern Nili Fossae

[Mangold et al., 2007; Mustard et al., 2007; Mustard et al., 2008]. From 72°E to 80°E, smectite-bearing basement rock is well-exposed by erosion in numerous locations, including in a 600-m thick section of the walls of the largest Nili Fossae trough. As discussed further by Mustard et al. [this issue], this Fe/Mg smectite-bearing unit may represent megabreccia constituting the early Noachian Martian crust. It is composed of variably altered blocks bearing Fe/Mg smectite and low-calcium pyroxene in a smectite- bearing matrix. When eroded, the brecciated unit commonly exhibits linear ridges, which may represent former conduits of fluid flow. Some Fe/Mg smectite-bearing materials have also undergone sedimentary transport and are a component in layered deposits cropping out in mesas in northern Nilo-Syrtis chaos, filling the fossae, filling impact craters, and forming deltaic deposits in Jezero crater [Mustard et al., this issue; Ehlmann et al., 2008b]. 77

The distribution of kaolinite is mostly restricted to the eastern portion of the study area, in and around the Nili Fossae. Kaolinite-bearing units occur as a distinctively bright-toned layer, always on top of the Fe/Mg smectite-bearing units. The kaolinite- bearing layer overlies both types of Fe/Mg smectite-bearing deposits—massive crustal smectites as well as layered sedimentary smectites. West of the Nili Fossae, Figure 13 shows a thin (10s of meters thick at most) layer of kaolinite-bearing material being exposed by the erosion of an overlying cap rock unit that lacks a mineralogically distinct spectral signature in CRISM data but is the highest stratigraphic unit in and around the

Nili Fossae [Mustard et al., this issue]. The kaolinite-bearing unit overlies a much thicker, ridged Fe/Mg smectite-bearing unit. Hundreds of miles to the east, kaolinite-bearing units also overlie massive crustal phyllosilicates. Figure 14 shows a kaolinite-bearing unit overlying a smectite-bearing unit along the wall of Nili Fossae trough, a 5 km long portion of which has been faulted and forms a down-dropped slump block. The kaolinite occurs in a bright toned unit, well-exposed beneath a mostly removed thin mantle. The kaolinite is exposed in terrain elevated ~80 m relative to its surroundings in MOLA DEM data, although without a higher resolution topographic model for the surface it is difficult to determine the dip and thickness of the kaolinite-bearing unit; 80 m should be considered an upper bound. The kaolinite-smectite stratigraphy is present even on the downdrop block, indicating kaolinite formation likely predates the fossae opening or at least the last significant tectonic activity.

Just to the northwest of this scarp is a 40 km crater that is over 1 km shallower than a nearby counterpart of the same diameter due to significant sedimentary fill (Figure

15). A CRISM image acquired over an eroded pit in the center of the crater shows that 78 although most of the fill is Fe/Mg smectite-bearing (A053, Figure 4a), the uppermost ~20 m of the exposed stack of fill has the distinctive spectral signature of kaolinite (Figure

15c). Interestingly, as in observations of kaolinite over massive crustal Fe/Mg smectite- bearing terrains, no bedding can be discerned in the kaolinite-bearing unit. The kaolinite unit in some cases does have a fractured appearance, as in Figure 15f, where the fractures appear to extend into an overlying thin mantle. However, along with its stratigraphic position, the unit’s most characteristic trait is that the kaolinite-bearing materials are typically 5-15% brighter than surrounding materials in both infrared albedo at 1.3 μm

(IRA parameter; Pelkey et al., 2007) and reflectance at 0.77 μm (R770).

In the easternmost portion of Nili Fossae, 76°E to 80°E, the occurrence of kaolinite is less common. Instead, beneath a capping unit but above Fe/Mg smectites are olivine-bearing rocks [Mustard et al., 2008; this issue] that have in places been altered, in some cases to magnesium carbonate [Ehlmann et al., 2008a]. To date, only one imaged location contains both carbonate- and kaolinite-bearing units in a distinctive stratigraphy.

In this image, it appears that a bright kaolinite-bearing unit overlies a polygonally fractured carbonate-bearing rock unit (Figure 16), which is partially obscured by olivine sands derived from the regional olivine bedrock unit [Hamilton and Christensen, 2005;

Mustard et al., 2008]. Further imagery to assess the kaolinite-carbonate relationship will be sought.

As discussed by Mustard et al., [this issue] the olivine-bearing rocks in eastern

Nili Fossae drape pre-existing topography and fill in topographic lows. Figure 17 shows the typical three-unit stratigraphy observed: (1) Fe/Mg smectite-bearing units are overlain by (2) olivine-bearing rocks, which in some places are altered and contain Mg carbonate, 79 and are in turn overlain by (3) a cap rock which lacks a mineralogically distinctive signature in CRISM data. In this particular location, a crater was formed in Fe/Mg smectite-bearing units, subsequently filled and covered by the olivine-bearing and cap rock units, and then eroded by surface fluvial activity that formed a valley draining into one of the fossae to the southeast. Exposed, ridged Fe/Mg smectite megabreccia is capped by mesas of banded olivine overlain by a cap rock without a distinctive mineralogic signature (Figure 17). Around the margins of the cap, where water is likely to have flowed, these units have in some places altered to Mg carbonate (green, Figure

17c). However, some of the olivine shows signs of alteration of a different nature

(magenta in Figure 17c-e) with an absorption centered at 2.32 m. This absorption occurs at a wavelength too long to be Fe/Mg smectite, which has its absorption at 2.30-2.31 m

(3E12, Figure 4a) and lacks the 2.5-μm absorption, which is characteristic of carbonate

(Figure 17d). The band centers of these spectra are more similar to saponite or to the serpentine observed elsewhere in eastern Nili Fossae. To date, serpentine has only been definitively identified in two images in very small exposures in and around the fossae

(Figure 2, 6). One is in FRT0000B8C2 where it seems contiguous with olivine and Mg carbonate-bearing rocks. The second, in heavily eroded terrain in FRT0000ABCB does not correlate with a distinctive unit in a stratigraphy.

Immediately around the Nili Fossae, alteration phases are also sometimes associated with fluvial features. Mg carbonate and Fe/Mg smectite are found in the valleys feeding into Jezero crater and in transported sediments within the Jezero crater deltas. Jezero crater is a 45-km crater, filled with sedimentary deposits and fed by a

15,000 km2 catchment, which is thought to have hosted a paleolake (Fassett and Head, 80

2005; Ehlmann et al., 2008b). Figure 18 shows the relative distribution of carbonate vs. smectite near Jezero crater’s western delta, which extends outward from a valley to the west (not pictured the image). Spectra from light-toned units in the well-formed western delta appear to indicate packages of distinctive composition, some more Mg carbonate enriched and another more Fe/Mg smectite enriched (Figure 18b, c). The clearest carbonate signatures—those least contaminated by smectite—found to date are from

HRL000040FF and rocks draped on the western wall of Jezero crater that underlie the western delta (Figure 12). The northern delta appears somewhat more Fe/Mg smectite enriched than the western delta, although all sediments probably contain both alteration minerals. Magnesium carbonate is also found near an eroded fluvial channel in the Nili

Fossae (green in Figure 17a). This same channel (arrow to cyan, Figure 17a) hosts the hydrated mixed phase (silica+zeolite or silica+sulfate; Figure 11).

5.3 Western cratered province

In the far west of the study area, along the eastern margin of the Antoniadi basin

(62° to 67°E), a distinctly different suite of alteration minerals occurs, accompanying iron-magnesium smectite (Figure 2). Chlorite is typically found as is hydrated silica. In two craters, the distinctive spectral signature of analcime has also been observed.

In this area, the Fe/Mg smectite and chlorite are found in bedrock units. In four

CRISM images acquired within the Antoniadi basin, chlorite and smectite occur in knobs and ridges extending over several kilometers, which have been embayed by younger units with no distinctive mineralogic signature in NIR data. In two 25 km impact craters, chlorite and smectite are found in the central peak. The northernmost crater shows the brecciated nature of the central peak (Figure 19) with 100 m angular boulders set in a 81 rock matrix. The mineralogic signature of smectite and chlorite appears in both blocks and matrix material. Whether spectra appear more chlorite-like or Fe/Mg smectite-like varies spatially but does not appear to correspond to unique geomorphic units. Within the northern 25 km crater, chlorite/smectite is also found in rocks exposed on the southwestern crater wall (Figure 19b). Although most chlorite-bearing materials in the region around the Nili Fossae also have a 1.48-μm band indicative of the presence of prehnite, the chlorite in these craters lacks this prehnite spectral feature (Figure 5b).

Within the two 25 km craters, the sodium zeolite analcime (Figure 9) is found in rocks in and around the central peak (Figure 19). The size of zeolite-bearing outcrops is comparatively small, no more than a few tens of CRISM pixels. In some cases, the physical properties of zeolite-bearing rocks are distinctive. Figure 19d shows analcime in brighter-toned rocks being exhumed from beneath a smooth covering material that lacks a mineralogically distinctive signature. In other places, analcime is found associated with large boulders at the base of the central peak.

The occurrence of hydrated silica is common throughout this part of the study area; however, to date this phase has not been found in any rock unit. Rather, the hydrated silica is found in mobile materials. In the northern 25 km crater, hydrated silica occurs in a ring around the central peak, which HiRISE imaging shows is a mobile unit of aeolian dunes and ripples (Figure 19c; 9312 in Figure 10). Smaller patches of hydrated silica-bearing sands are located in local depressions and along scree slopes on the southern wall of the 25 km crater and within the Antoniaidi basin. Interestingly, the locations of the hydrated silica correspond to the locations identified by Bandfield [2006] using Thermal Emission Spectrometer (TES) and THEMIS data, as units of granitoid or 82 quartzofeldspathic composition (Figure 20). Materials ringing the central peak mapped as quartzofeldspathic [Bandfield et al., 2004] correspond to the aeolian materials mapped with CRISM as hydrated silica.

5.4 Central province: North Syrtis craters

In craters located near the margins of the Syrtis Major flow immediately west and south of Nili Fossae, chlorite and prehnite are the areally dominant alteration minerals, and Fe/Mg smectites are considerably less prevalent. In contrast to the craters in the

Western province, the less areally abundant phases present within each of these craters vary considerably among craters. The northernmost of the craters has associated K mica while the southernmost craters have hydrated silica or, outside of the crater, kaolinite

(Figure 2).

Figure 21 shows a 50 km diameter crater at the northern margin of the Syrtis

Major lava flows. High resolution images indicate that the crater of interest was formed after emplacement of the Hesperian Syrtis Major formation; its ejecta overlie unaltered

Syrtis Major lavas that fill a crater to the southeast. The crater of interest either formed by an impact into the side of a hill or has been heavily eroded on one side. MOLA data show the eastern side is nearly 1 km lower than the western side (Figure 21c). CRISM images have been acquired of the crater’s wallrock and ejecta. All scenes display some evidence for hydrated silicates (Figure 21b). Whereas the wall rocks contain phyllosilicates in only small, isolated knobs, spectral signatures of alteration minerals are found over nearly the entire surface in scenes acquired along the eastern wall and in terrains further east.

Prehnite and chlorite (454E, Figure 5b) are the dominant minerals, displayed as green to yellow—yellow indicates a more prominent 2.35-μm absorption—in Figure 21d-e. As 83 discussed in section 4.2, these two minerals are difficult to distinguish in VNIR data. The prehnite/chlorite appears to originate from rocky knobs and small impact craters and is then dispersed by erosion. A few knobs have materials bearing the spectral signature of K mica (Figure 8) although these do not appear to be geomorphically distinct from prehnite/chlorite-bearing knobs (Figure 21d,e). Near some of the small craters, polygons with sides tens of meters in length have been observed, and these terrains are also chlorite/prehnite-bearing (Figure 21f).

Immediately southwest of Nili Fossae, prehnite/chlorite occurs in wall rocks (e.g.

50F2, Figure 5b) and in the central peak materials of two 40 km impact craters (Figure

22). In their ejecta, Fe/Mg smectite is most common but chlorite/prehnite is found in a few small blocks. In the northern crater, small outcrops of hydrated silica (Figure 10) are found within the crater. A mélange of mixed hydrated silica, chlorite, and prehnite- bearing materials are found in the central peak, and the parameter maps as well as their spectra show these are mixed at 18 m/pixel CRISM resolution (Figure 22). In contrast, no alteration minerals have been found in the interior of the southern crater, which has been covered by Syrtis Major lava flows. Around this crater’s exterior are numerous small outcrops bearing kaolinite. Small craters that post-date the larger 40 km crater expose

Fe/Mg smectites and chlorite in their walls and ejecta (Figure 22).

6.0 Discussion

The diversity of alteration minerals and, moreover, their distinctive associations and geomorphic expressions suggests that the nature of aqueous activity varied in space and time across the greater Nili Fossae region. After briefly discussing possible scenarios 84 for formation of the Fe/Mg smectites, which are observed throughout the study area, we focus on key findings specific to the region: (1) in and around the Nili Fossae, evidence for multiple episodes of aqueous activity with neutral to alkaline waters during the period from the early Noachian to the late Hesperian and (2) in cratered terrains, evidence for low temperature metamorphic or hydrothermal aqueous alteration. We then discuss some important outstanding questions that could form the basis for future work.

6.1 Widespread Fe/Mg smectites

As is the case planet-wide, the high resolution view provided by CRISM targeted images at Nili Fossae has greatly expanded the number of terrains known to contain

Fe/Mg smectite [Mustard et al., 2008]. In the greater Nili Fossae region, Fe/Mg smectites are found in bedrock as the lowermost stratigraphic unit in images sampling over 100,000 km2. The spatial extent of the Fe/Mg smectite seems to exclude local lacustrine or volcanic hydrothermal processes as the primary formation mechanism. Rather for such widespread distribution, a more widespread process such as pedogenic leaching during alteration at the surface or subsurface hydrothermal activity related to crustal cooling or impacts is required. The brecciated nature of much of the smectite-bearing crust suggests the primacy of the role of impacts in distributing Fe/Mg smectites, if not also in creating them.

Had there been insufficient time for reaction kinetics to produce thermodynamically favored phyllosilicate products, amorphous weathering products (e.g. allophane, palagonite or their precursors) would be found rather than the spectrally distinct phyllosilicates that we observe. Hence, based on terrestrial analogs, the materials in the Fe/Mg smectite-rich basement probably altered over at least thousands of years 85

[Eberl, 1984; Price et al., 2005], although the abundance of Fe/Mg smectite constituents relative to unaltered parent materials is an important additional constraint on the duration of aqueous activity. Estimates at Nili Fossae from phyllosilicate-bearing terrains are 20-

35% phyllosilicate, a figure that encompasses both abundance in bedrock and any covering sand or dust [Poulet et al., 2008]. Nontronite and saponite typically form at low temperatures (<200°C) but can form and persist—that is, not undergo transformation to other minerals—at temperatures up to 600°C under the right fluid chemistries so their thermal stability does not by itself constrain the likely formation environment [Meunier,

2005]. The partially altered Fe/Mg smectite-bearing brecciated rock unit that comprises the >600m thick basement unit exposed by the Nili Fossae is representative of deep phyllosilicates of the early Noachian crust [Mustrad et al., this issue; Murchie et al., this issue]. It is exposed by extensive erosion around the fossae, but exhumed only locally by impact craters farther to the south and west. The spatially widespread nature of this

Fe/Mg smectite suggests long-lived (likely at least ~103 years), pervasive water-rock interaction that was a regionally or globally important process for altering the earliest

Martian crust.

6.2 Multiple episodes of aqueous activity around Nili Fossae

Although smectites can form in numerous environmental settings, their accompaniment by additional alteration minerals provides additional constraints on timing, temperature, pressure, and aqueous geochemistry during alteration. In fact, the

Nili Fossae region shows evidence for three episodes of aqueous activity: (1) alteration forming Fe/Mg smectites, (2) alteration to form overlying kaolinite or magnesium 86 carbonate, and (3) erosion of units bearing these alteration minerals by fluvial activity and sapping channel formation.

The coherent stratigraphy of layered phyllosilicates, i.e. kaolinite or carbonate bedrock always overlying smectite, in the immediate vicinity of Nili Fossae would not have been preserved during excavation and ejecta emplacement from the formation of the

Isidis basin. Hence, a constraint is provided on (2), the timing of the formation of kaolinite and magnesium carbonate in eastern Nili Fossae; it postdates the Isidis impact, which is dated at ~3.96 Ga or approximately Mid-Noachian [Werner, 2005]. A control on the distribution of carbonate versus kaolinite appears to be parent rock composition. The eastern portion of Nili Fossae is the site of the largest exposure of olivine-rich rocks on

Mars [e.g. Hoefen et al., 2003; Hamilton and Christensen, 2005; Poulet et al., 2007;

Koeppen and Hamilton, 2008] (Figure 1). These precursor ultramafic rocks in eastern

Nili Fossae have, in some places, partially altered to Mg carbonate [Ehlmann et al.,

2008a]. In areas near the western part of Nili Fossae, where olivine is not usually identified in NIR spectral data, kaolinite instead typically occupies the middle stratigraphic layer between the Fe/Mg smectite and a capping rock.

To the best resolution of HiRISE (25 cm/pixel), the tens of meters thick kaolinite- bearing unit is not bedded as might result from sedimentary transport. This may indicate that the kaolinite formed in place by alteration of parent rocks just as the magnesium carbonate apparently formed in place by alteration of olivine-rich rocks [Ehlmann et al.,

2008a]. A plausible formation mechanism for a kaolinite layer overlying smectite deposits is pedogenic-type leaching of pre-existing phyllosilicates, leading to loss of Ca,

Mg, and Fe ions from Fe/Mg smectite and its transformation to kaolinite. For example, 87 kaolin-bearing soils form from basalts in Hawaii under intermediate leaching conditions; smectites are found in drier areas with less throughput of water whereas gibbsite and iron oxides are found in wetter environments [Bates, 1962; Eberl, 1984]. Alternating wet-dry cycles with prolonged periods of aridity appear to promote halloysite development

[Ziegler et al., 2003] on basalt, even under conditions that thermodynamically favor smectite. At Nili Fossae, the kaolin-bearing weathered soil unit might have subsequently been covered by the cap rock material that presently overlies it. Kaolinite can, however, also form hydrothermally via precipitation from fluids. If a hot impact melt sheet or hot ejecta were emplaced on water-bearing smectite-rich rocks, circulation of hydrothermal fluids might lead to kaolinite formation in the zone along the contact between smectite and ejected Isidis materials, which may constitute the cap rock. While the morphology and stratigraphy of the cap rock unit is well-characterized [Mustard et al., this issue], its composition and method of emplacement have not yet been ascertained. From spectral data, we see no evidence for the high temperature polymorphs dickite and nacrite, which can be distinguished spectrally from kaolinite; however, kaolinite can also form and persist at high temperatures [Zotov et al., 1998; Fialips et al., 2003]. As discussed in section 4.4, the spectral data are consistent with kaolinite, halloysite or a mixture of kaolinite and smecitite (mixed layer clay or physical mixture) (Figure 7) any of which are consistent with kaolinite formed in pedogenic processes.

Like kaolinite, magnesium carbonates also can form in both surface and subsurface environments. On Earth, magnesite is uncommon but where found in large deposits is typically formed in playa environments fed by ultramafic catchments, during near surface alteration of serpentinized rock bodies, or from hydrothermal or 88 serpentinization reactions at elevated temperatures by fluids circulating through ultramafic rocks in the subsurface [Moeller, 1989]. If in fact the serpentine found in Nili

Fossae is associated with olivine-carbonate rocks as suggested by the band position of the

2.32 m absorption in some exposures of partially altered olivine (Figure 17), one of the latter two mechanisms involving serpentinization is favored. Additional images of serpentine and partially altered olivine deposits will be acquired to assess the relationship between the two minerals.

Sapping channels and valleys [Mangold et al., 2007] cut the layered phyllosilicates in Nili Fossae. For example, layered Fe/Mg smectite and kaolinite sediments are cut by a channel draining a crater (Figure 15), and fossae walls have also been significantly eroded by alluvial activity (e.g. Figure 17) [Mustard et al., this issue].

Mangold et al. (2007) constrain the timing of this erosive, fluvial episode (3) to be Late

Noachian to Hesperian, which effectively brackets the timing of the formation of the kaolinite and carbonate (2). In most locations, this last episode of aqueous activity did not result in formation of alteration materials in transported sediments or in source watersheds [Mangold et al., 2007]. However, in a few places (e.g Figure 11, 17a), hydrated minerals are precipitated in channel systems associated with this activity. This deposit appears to be a mixture of hydrated silica with either a sulfate or zeolite.

Collectively, the alteration minerals in Nili Fossae point to waters of neutral to alkaline pH, which contrasts to acidic pH conditions indicated in other locations and perhaps in later time periods [Hurowitz and McLennan, 2007]. Carbonate persists to the present, associated with Hesperian channel systems. This indicates that waters were probably not acidic or that acidic waters were not in contact with the carbonate-bearing 89 rocks for very long because this would have dissolved the carbonate [Ehlmann et al.,

2008a]. In fact, carbonate is an important component of Fe/Mg smectite-bearing deltaic sediments within Jezero crater. The similar layering of clays—Al-rich phyllosilicates over Fe/Mg smectite—at multiple locations across Mars [e.g. Bishop et al., 2008;

McKeown et al., this issue; Noe Dobrea et al., this issue; Murchie et al., this issue] may indicate similar formational processes. However, unlike the kaolinite seen in Mawrth

Vallis [McKeown et al., this issue], in the layered clays at Nili Fossae there is no evidence for any additional associated phase such as hydrated silica that would broaden the absorption at 2.2 m. Co-precipitation of amorphous silica and kaolinite occurs for low pH but is less common at mildly acidic to neutral pH [Fialips, et al., 2000]. The observation of kaolinite-bearing rocks apparently in direct contact with carbonate-bearing rocks in at least one location (Fig. N) also suggests that this kaolinite did not form under acidic conditions. Rocks in and around Nili Fossae thus likely preserve a record for multiple episodes of aqueous alteration of a neutral to alkaline character, different from many sites elsewhere on Mars.

6.3 Mineralogic evidence for low temperature metamorphic or hydrothermal alteration

In cratered terrain surrounding the fossae (Noachian etched plains (Nple) [Greeley and Guest, 1987]), the types of alteration minerals are distinctly different from those in the immediate vicinity of the fossae. Chlorite, prehnite, hydrated silica, and analcime do not occur in a coherent stratigraphy but rather are exposed in breccia blocks, isolated outcrops or in eroded debris. However, in spite of the lack of clear geologic context, some of these minerals—particularly when considering the mineral associations in which 90 they are found—point to aqueous alteration that occurred at elevated temperatures in the subsurface.

The 1.48 m band indicating the existence of prehnite (Figure 5) in mixtures in numerous craters in Nili Fossae provides the clearest evidence for hydrothermal or low temperature metamorphic activity. Prehnite forms only under a highly restricted set of conditions: temperatures of 200-350ºC, pressure < 3 kbar, and a low partial pressure of

CO2 (e.g. XCO2 < 0.004 at 1.5 kbar) [Schiffman and Day, 1999; Robinson and Bevins,

1999]. On Earth, it is typically found in association with pumpellyite and chlorite in low temperature metamorphosed rocks or in amygdules precipitated in the vesicles of basalt.

Poulet et al. [2008] have reported the presence of pumpellyite within this crater based on results from spectral mixture modeling. The prehnite now exposed in the central peak of the crater in north Syrtis Major (Figure 22) could not have formed in its current setting from waters in contact with the CO2-rich atmosphere. Rather, it probably formed from aqueous alteration at elevated temperatures in the subsurface and was brought to the surface by the impact. An alternative, less plausible scenario is formation of prehnite post-impact during hydrothermal circulation of waters in the central peak, which were out of contact with the atmosphere, and then subsequent extensive erosion of the central peak to expose the prehnite. However, the existence of prehnite/chlorite in both the central peak and crater ejecta suggests it likely formed during subsurface alteration that pre-dates the impact. As reported by Buczkowski et al. [this issue], occurrences of prehnite associated with craters around the Argyre basin suggest that this process was not unique to Nili Fossae. 91

The small K mica-bearing rock outcrops among prehnite- and chlorite-bearing ridges are puzzling geologically. Though it has been found elsewhere in the southern highlands [Fraeman et al., 2009], so far, K mica in Nili Fossae is restricted to one 50 km crater (Figure 2, 21). It is difficult to tell whether the K mica, chlorite, and prehnite- bearing rocks constitute the ejecta of the crater or whether the entire eastern side of the crater has been heavily eroded and the chlorite, prehnite, and illite simply represent exposed bedrock. In either case, the eastern side of the crater is some 1 km lower than the western side so rocks from deeper in the crust are probably being sampled. On Earth, K micas, e.g. muscovite, can be primary in granitic terrains. However, on Mars where basalt is dominant, K mica more likely results from hydrothermal activity or diagenesis. A key diagenetic process forming illite on Earth is conversion of smectite to illite upon burial, a process that is time and depth dependent and usually begins at temperatures of 50-80ºC

[Hower, 1976]. Dioctahedral smectites are converted first to mixed layer illite-smectites and then to illite of various polytypes. Higher grade metamorphic/hydrothermal activity can generate muscovite. For trioctahedral smectites, with increasing temperature the reaction proceeds as saponite to corrensite (intermediate chlorite-smectite) and to chlorite

[Merriman and Peacor, 1999]. In weakly metamorphosed rocks and sediments with a high ferromagnesian component, e.g. Fe/Mg smectites, a chlorite-illite assemblage with chlorite dominating can occur and may explain the alteration minerals around the 50 km crater.

In the far west of the study area, craters have associated analcime-bearing rocks and aeolian hydrated silica deposits. Analcime and hydrated silica can originate in weathering reactions or from hydrothermal activity, and zeolites have been suggested as a 92 possible constituent of Mars dust [Ruff, 2004]. In particular, the most common environments for analcime are highly alkaline lake waters [Eugster, 1980], weathering of tephra [Sheppard and Hay, 2001], and hydrothermal circulation <~200ºC as has been extensively studied for Icelandic basalts and Massif Central in France [e.g. Weisenberger and Selbekk, 2008; Robert and Goffe, 1993]. No evidence for layered sediments, which might indicate one of the first two mechanisms, exists in the two 25 km craters near the

Antoniadi basin. The craters have been intensively studied in both near-infrared and thermal wavelengths to determine the composition of rocks and aeolian materials. TES deconvolution results show the central peak spectral unit to be enriched in quartz (total,

7%), feldspar (9%), and sheet silicates/high Si glass (33%) and depleted in pyroxene

(9%) relative to nearby Syrtis Major plains materials [Bandfield et al., 2004]. CRISM detects analcime, hydrated silica, chlorite, and Fe/Mg smectite. A mineral assemblage of

Fe/Mg-rich smectites or chlorites with accessory zeolite, silica, quartz, and K-feldspar results from hydrothermal alteration in terrestrial craters [e.g. Allen et al., 1982; Naumov,

2005] and suggests a possible reinterpretation of the “granitoid” material in Bandfield

[2006] as hydrothermal in origin rather than igneous, as has been proposed for other

TES-detected high silica materials near Hellas [Bandfield, 2008]. The moderate TES- modeled abundances of quartz and K-feldspar and their association with CRISM-detected hydrated Si-OH bearing phases (e.g. hydrated silica), phyllosilicates, and zeolites are consistent with such a formation process. Quartz, K-spar, analcime, and amorphous silica can form as amgydules in vesicles. Upon erosion of these rocks, more resistant silica-rich phases may be left behind as an aeolian lag, explaining the association of hydrated silica with mobile materials. 93

Variations in the temperature, water-rock ratio, and fluid chemistry of hydrothermal systems may be responsible for these distinctive mineral assemblages observed in Nili Fossae region craters. In all cases, however, fairly low temperature

(<350ºC) metamorphic and/or hydrothermal processes are indicated by the mineral phases so far observed.

6.4 Key questions for future work

The diversity of alteration minerals detected in the region around the Nili Fossae is the highest yet observed on Mars. Why is the diversity so great here—is it simply a result of a single event, the Isidis impact, and aqueous alteration, e.g. hydrothermal, immediately following and related to basin formation? The original Isidis crater has been heavily modified by tectonism, erosion, and deposition to create the Isidis basin seen today, and it is not entirely clear how far into the original impact basin the Nili Fossae are found. The fossae likely correspond to a megaterrace formed between the inner and outer basin rings [Mustard et al., 2007]. The brecciated Fe/Mg smectites in the lowermost stratigraphic layer are the impacted materials, and olivine rich rocks at the base of an otherwise mineralogically indistinct cap unit are likely the product of an impact melt sheet [Mustard et al., this issue]. The minerals found in smaller craters west of the basin would, if directly related to the Isidis impact, have possibly been generated during hydrothermal activity within and around the outer basin ring. This is plausible given computational models of hydrothermal cooling that indicate hydrothermal systems can endure for 103 years or longer with increasing duration for larger impacts [Rathbun and

Squyres, 2002; Abramov and Kring, 2005]. However, the coherency over 100’s of km of the layered stratigraphy in the eastern part of the study region is less easily explained by 94 impact-induced processes. Moreover, the timing of at least the latest fluvial activity definitively post-dates opening of the fossae [Mangold et al., 2007]. These both suggest that any hydrothermal processes that might have been related to the Isidis impact represent only one of several episodes of regional aqueous activity.

In terms of understanding the extent and duration of alteration in various terrains, initial efforts [Poulet et al., 2008] to calculate modal mineralogy in phyllosilicate-bearing terrains should be expanded. An essential part of this effort will be expanding the number of well-characterized minerals and rock samples in publicly-available spectral libraries as well as making laboratory measurements to measure their optical constants.

While the mineralogic evidence points to low temperature metamorphic or hydrothermal processes south and west of Nili Fossae near craters, a key question is their timing. Did these processes pre-date or post-date the impact craters with which they are associated? Understanding this requires high-resolution models of impact cratering processes. Some models predict shock temperatures near central peaks would be too high for phyllosilicates or silicate hydroxides to survive [Marzo et al., 2008; Fairen et al.,

2008]; others have lower initial temperatures [Barnhart et al., 2008]. The ability to model discrete packets of materials within the pre-impact surface and their movement and temperature throughout compaction and compression, excavation, and adjustment stages is key to understanding the degree of heterogeneity of temperatures experienced by impacted materials as well as where they end up after the most intense period of the modification stage. This would allow understanding the likely distribution and thermal history of impacted subsurface units. Also important will be moving beyond which minerals are thermodynamically stable to understanding the kinetic rates for formation 95 and destruction of alteration minerals [Schwenzer and Kring, 2009]. By coupling detailed knowledge of temperature and mass movement with understanding of the rates of mineral formation/destruction, the question of pre- or post-impact alteration can be rigorously assessed. Also important will be advances in interpretation of VNIR spectra, in particular of mixed layer clays such as illite-smectite and smectite-chlorite, in order to distinguish the degree of diagenesis/metamorphism

As CRISM continues to map the surface of Mars, the uniqueness of the mineralogic diversity found in the Nili Fossae region will be assessed. Further efforts will also examine the S detector data to look for evidence of aqueous alteration provided by iron oxide minerals. Was the region around the Nili Fossae unusually wet in comparison to other regions of Mars? How much of the Noachian megabreccia [McEwen et al., 2008] of Mars is altered? Already, some of the alteration minerals—K mica, hydrated silica, chlorite—have been found elsewhere in the southern highlands, associated with craters in the Terra Tyrrhena region [Fraeman et al., 2009]. Diverse alteration minerals have been mapped first in the well-exposed Nili Fossae region, following up on regional scale

OMEGA phyllosilicate detections; however, with future exploration, we may find aqueous activity, including low grade metamorphic or hydrothermal activity, was much more widespread on ancient Mars than previously known.

7.0 Conclusions

In the Noachian terrain around the Nili Fossae, alteration minerals identified by

CRISM based on distinctive absorptions from 1.0-2.6 μm include nontronite and magnesium smectites, chlorite, prehnite, serpentine, kaolinite, K mica, analcime, 96 hydrated silica, and magnesium carbonate. While Fe/Mg smectites are found throughout the study area over a much greater area than previously known, the distribution of the other discovered alteration minerals is not homogeneous. Rather, these tend to occur in provinces with distinctive assemblages of alteration minerals and geomorphic settings.

Stratigraphic relationships among units bearing alteration minerals can be discerned in the eastern part of the study area, in eroded terrains in and around the fossae, where layered clays with Fe/Mg smectite, Mg carbonate, and kaolinite are found. In bedrock units, Mg carbonate and kaolinite occur above Fe/Mg smectite-bearing units.

Magnesium carbonate is more common near the eastern Nili Fossae and kaolinite is more common in the west. This probably reflects the nature of the parent material, namely the prevalence of more olivine-rich rocks in the east that can alter to magnesium carbonate.

In and around impact craters west and south of the Nili Fossae, illite, analcime, and prehnite serve as mineralogic indicators for alteration at elevated temperatures in the subsurface. Low temperature metamorphism or hydrothermal activity is most definitively indicated by prehnite where alteration occurred out of contact with the CO2 atmosphere and at temperatures between 200-350°C. While Fe/Mg smectite is still found associated with these impact craters, chlorite- and prehnite-bearing materials are much more common in the cratered terrains west and south of the Nili Fossae and may reflect conversion of smectites to chlorite by burial and/or exposure to higher temperatures at depth.

Key questions for future work include constraining whether aqueous alteration at elevated temperature pre-dated or post-dated impact structures and more definitively constraining the environment of carbonate and kaolinite formation. Regardless, the 97 mineralogic diversity of the region around the Nili Fossae indicates multiple episodes of aqueous activity in multiple distinct environments. An initial phase of Fe/Mg smectite formation in the early- to mid-Noachian was apparently followed by episodes of subsurface or surface aqueous mineralogic alteration, culminating finally in Hesperian fluvial reworking that led to little mineralogic alteration. The resultant alteration minerals at Nili Fossae are more typical of those resulting from neutral to alkaline conditions rather than acidic conditions thought to have dominated other regions of Mars. The excellent exposure of bedrock units and the apparently long and complex history of aqueous activity in the region around the Nili Fossae make this part of Mars a compelling destination for future in-situ investigations.

Appendix Table 1. Full image IDs, locations, and number of pixels used for the numerators and denominators of CRISM spectra shown in Figures 4-12. All pixel locations are given for unprojected CRISM image data. Label in Figure CRISM Image Center Pixel(s) Number of ID pixels

Figure 4 9312 c/m (numerator) FRT00009312 X:337 Y:230 ROI, 104 9312 c/m (denominator) FRT00009312 X:337 Y:230, X:334, ROIs, 112 Y:155 A053 r/g (numerator) FRT0000A053 X:457 Y:36 ROI, 94 A053 r/g (denominator) FRT0000A053 X:457 Y:141 ROI, 71 9D44 c/m (numerator) FRT00009D44 X:207 Y:181 ROI, 116 9D44 c/m (denominator) FRT00009D44 X:207 Y:181 ROI, 99 3E12 r/g (numerator) FRT00003E12 X:127 Y:397 ROI, 295 3E12 r/g (denominator) FRT00003E12 X:127 Y:198 ROI, 386 A053 b/y (numerator) FRT0000A053 X:274 Y:78 ROI, 205 A053 b/y (denominator) FRT0000A053 X:274 Y:154, X:264 Y:170 ROIs, 176 9D44 r/g (numerator) FRT00009D44 X:561 Y:143 ROI, 28 9D44 r/g (denominator) FRT00009D44 X:561 Y:425 ROI, 19 64D9 r/g (numerator) FRT000064D9 X:459 Y:7 ROI, 44 64D9 r/g (denominator) FRT000064D9 X:459 Y:25 ROI, 28 527D r/g (numerator) FRT0000527D X:88 Y:103 ROI, 46 527D r/g (denominator) FRT0000527D X:88 Y:252 ROI, 81 50F2 c/m (numerator) FRT000050F2 X:74 Y: 361 ROI, 1419 50F2 c/m (denominator) FRT000050F2 X:74 Y: 294 ROI, 1172

Figure 5 50F2 s1a FRT000050F2 X:125 Y:434 5x5 50F2 s2a FRT000050F2 X:125 Y:310 5x5 50F2 s1b FRT000050F2 X: 216 Y:66, X:283 Y:82 5x5+5x5 50F2 s2b FRT000050F2 X: 216 Y:177, X:283 Y:12 5x5+5x5 454E s1 FRT0000454E X:243 Y:382 5x5 454E s2 FRT0000454E X:243 Y:26 5x5

Figure 6 ABCB s1 FRT0000ABCB X:253 Y:60 5x3 ABCB s1 FRT0000ABCB X:253 Y:15 5x3 B8C2 s1 HRL0000B8C2 X: 70 Y: 174 7x7 B8C2 s2 HRL0000B8C2 X: 70 Y: 59 7x7

Figure 7 ABCB s1 FRT0000ABCB X:387 Y:222 3x3 ABCB s2 FRT0000ABCB X:387 Y:338 3x3

Figure 8 2FC5 s1 HRS00002FC5 X:81 Y:14 2x2 2FC5 s2 HRS00002FC5 X:81 Y:53 2x2 454E s1 FRT0000454E X:100 Y:405 5x5 454E s2 FRT0000454E X:100 Y:373 5x5

Figure 9 9312 s1 FRT00009312 X:174 Y:400 3x3 9312 s2 FRT00009312 X:174 Y:291 3x3

Figure 10 9312 s1 FRT00009312 X:154 Y:170 11x11 9312 s2 FRT00009312 X:154 Y:370 11x11 9EA3 s1 HRL00009EA3 X:157 Y:359 ROI, 71 9EA3 s2 HRL00009EA3 X:156 Y:370 ROI, 79 99

Figure 11 A4FC s1 FRT0000A4FC X:191 Y:122, X:113 ROIs, 136 Y:211, X:61 Y:187 A4FC s2 FRT0000A4FC X:191 Y:134, X:113 ROIs, 144 Y:178, X:61 Y:375

Figure 12 A09C s1 FRT0000A09C X:50 Y:98 ROIs, 277 A09C s2 FRT0000A09C X:50 Y:186 ROIs, 522 40FF s1 HRL000040FF X:42 Y:347, X:186 Y:367, ROIs, 1425 X:263 Y:296 40FF s2 HRL000040FF X:76 Y:238, X:85 Y:76, ROIs, 5277 X:237 Y:22, X:263 Y:91

Appendix Table 2. Source and full sample identifier for the laboratory library spectra shown in Figures 4-12. Label in Figure Source Source sample ID Notes

Figure 4 Fe-smectite SWa-1 Clark et al., 2007 Nontronite SWa-1.a Nontronite NG-1 Clark et al., 2007 Nontronite NG-1.a Hectorite SHCa-1 Clark et al., 2007 Hectorite SHCa-1 Sepiolite SepNev-1 Clark et al., 2007 Sepiolite SepNev-1 Saponite SapCa-1 Clark et al., 2007 Saponite SapCa-1

Figure 5 Prehnite Clark et al., 2007 Prehnite GDS613.a <60um Pumpellyite, Ca-rich RELAB ZE-EAC-001-c1ze01 * Pumpellyite, Mg-rich RELAB ZE-EAC-002-c1ze02 * Chlorite Clark et al., 2007 chlorite_smr13.4984 Clinochlore (Fe:Mg 0.57) Clark et al., 2007 clinochlore_gds158.5327 Clinochlore (Fe:Mg 1.5) Clark et al., 2007 clinochlore_gds157.5542

Figure 6 Antigorite Clark et al., 2007 Antigorite NMNH96917.b 165u Lizardite Clark et al., 2007 Lizardite NMHR4687 Saponite Clark et al., 2007 Saponite SapCa-1

Figure 7 Montmorillonite Clark et al., 2007 Montmorillonite SWy-1 Kaol-smect Clark et al., 2007 kaolsmect_h89fr2.25636 Kaolinite, wxl Clark et al., 2007 kaolinite_kga1.12117 Kaolinite, pxl Clark et al., 2007 kaolinite_pfn1_kga2.12176 Halloysite Clark et al., 2007 halloysite_nmnh106236.8988 Nacrite Clark et al., 2007 Nacrite GDS88 Dickite Clark et al., 2007 Dickite NMNH106242

Figure 8 Muscovite Clark et al., 2007 Muscovite GDS116 Tanzania Muscovite (hydrated) Clark et al., 2007 Muscovite GDS111 Guatemal Illite, 2M, hydT Clark et al., 2007 Illite IL101 (2M2) # Illite, 2M, sed. Clark et al., 2007 Illite GDS4 (Marblehead) # Illite, 1M, sed. Clark et al., 2007 Illite IMt-1.a # Illite, 1M, disordered Clark et al., 2007 Illite IL105 (1Md) #

Figure 9 Analcime Clark et al., 2007 Analcime GDS1 Zeolite Stilbite Clark et al., 2007 Stilbite GDS8 Zeolite 100

Mordenite Clark et al., 2007 Mordenite GDS18 Natrolite Clark et al., 2007 Natrolite HS169.3B Zeolite

Figure 10 Hydrated basaltic glass G. Swayze, pers. Hydrated Basaltic Glass MLUT03- shown at 10x comm. 9C 7torr 240K 45min actual reflectance Opal A/CT mixture Clark et al., 2007 Opal TM8896 (Hyalite)

Figure 11 Opal Clark et al., 2007 Opal TM8896 (Hyalite) Chalcedony Clark et al., 2007 Chalcedony CU91-6A Chabazite RELAB Zeolite Chabazite ZE-EAC-017 * LAZE17 Heulandite RELAB Zeolite Heulandite ZE-EAC- * 031LAZE31 Thomsonite RELAB Zeolite Thomsonite ZE-EAC- * 009LAZE09 Mg sulfate RELAB Magnesium Sulfate JB-JLB-366- 799F366 Fe sulfate RELAB Copiapite CC-JFM-013-A- F1CC13A

Figure 12 Magnesite CB-EAC-006A RELAB CB-EAC-006A-lacb06a * Magnesite CC-JFM-006B RELAB CC-JFM-006B-f1cc06b Siderite Clark et al., 2007 Siderite HS271.3B Coalingite RELAB CB-EAC-023-A-lacb23a * Mg carb. (hydromag.) RELAB JB-JLB-590-bkr1jb590 Magnesite+Hydromag. Clark et al., 2007 Magnesite+Hydroma HS47.3B Lizardite Clark et al., 2007 Lizardite NMNHR4687.b 165 Nontronite Clark et al., 2007 Nontronite NG-1.a Brucite Clark et al., 2007 Brucite HS34827.3B * sample descriptions by E. Cloutis available at psf.uwinnipeg.ca/Sample_Database.html # see also Kruse & Hauff, 1991 101

Acknowledgments. Mario Parente and Frank Seelos provided numerous useful suggestions on data analysis for this manuscript. Thanks to Lynn Carlson, Nicolas

Mangold, and Nancy McKeown for helpful advice on techniques for combining datasets to make geologic interpretations. We appreciate the many helpful discussions with members of the CRISM team and the ongoing, dedicated efforts of the CRISM-SOC to acquire this spectacular dataset. Reviewers Brad Dalton and Victoria Hamilton provided detailed comments that improved this manuscript. We also thank the CTX and HiRISE teams for their commitment to coordinated observations which enable high-resolution studies of Mars mineralogy with morphology.

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115 Figures

Figure 1. Thermal Emission Imaging System (THEMIS) daytime infrared mosaic with mafic mineralogic parameters overlain. In pyroxene-bearing terrains, the ratio of low-calcium pyroxene to total (low- and high- calcium) pyroxene is mapped with CRISM and OMEGA data using the Modified Guassian Model [mapping by P. Thollot according to the methods in Mustard et al., 2005] so that terrains enriched in low calcium pyroxene are green and those enriched in high calcium pyroxene are blue. The presence of olivine is indicated in red and mapped with OMEGA data using the olivine parameter [Pelkey et al., 2007]. 116

Figure 2. THEMIS daytime infrared mosaic of the study region west of Isidis indicating locations of hydrated minerals. Outlines of CRISM targeted images through August 31, 2008 (2008_244) are indicated by white outlines if no phyllosilicates are present and orange if Fe/Mg smectite is present in the observation. Presence of other hydrated minerals within the observation is indicated by the colored circles located next to the CRISM observation. Where observations have significant overlap, symbols indicating mineral presence are only shown once for the collection of observations. 117

Figure 3. THEMIS daytime infrared mosaic of the study region west of Isidis indicating four character hexadecimal IDs for images and spectra specifically referred to in this paper. Outlines of later figures are also shown. 118

Figure 4. CRISM spectra of iron magnesium smectite-bearing materials. (a) representative CRISM ratio spectra from numerous images throughout Nili Fossae (four character identifier provided; for full information, see Appendix and for location, see Figure 3) grouped by the frequency of occurrence for spectra with those characteristics (b) library spectra for iron magnesium smectites and sepiolite shown for comparison. Vertical lines at 1.4 and 2.3 μm indicate absorptions which shift in position as the proportion of iron vs. magnesium in smectite clays changes. The vertical line at 1.91 μm indicates the presence of

H2O and the 2.39-μm absorption indicates and OH combination tone. For details on spectra used, see Appendix. 119

Figure 5. CRISM spectra of chlorite and prehnite-bearing materials. (a) unratioed CRISM spectra from FRT000050F2 (15.0°N, 72.1°E) (b) ratioed CRISM spectra using numerators and denominators from (a) and for FRT0000454E (20.1°N, 70.1°E; numerator and denominator not shown). (c) prehnite, pumpellyite, and chlorites with varying iron content are shown for comparison. Vertical lines at 1.41, 1.48, 1.91, 2.01, 2.22, 2.28, and 2.35 μm indicate absorptions that show prehnite and chlorite are both indicated by CRISM spectra. For details on spectra used, see Appendix. 120

Figure 6. CRISM spectra of serpentine-bearing materials. (a) unratioed CRISM spectra from FRT0000ABCB (thin line; 20.9°N, 73.3°E) and HRL0000B8C2 (thick line; 17.3°N, 77.1°E). (b) ratioed CRISM spectra using numerators and denominators from (a). (c) the serpentines lizardite and antigorite and the Mg smectite saponite are shown for comparison. Vertical lines at 1.39, 2.11, 2.32, and 2.52 μm indicate absorptions that are together are diagnostic of the presence of serpentine. For details on spectra used, see Appendix. 121

Figure 7. CRISM spectra of kaolinite-bearing materials. (a) unratioed CRISM spectra from FRT0000ABCB (20.9°N, 73.3°E) (b) ratioed CRISM spectrum using numerators and denominators from (a). (c) Library spectra of kaolinite group minerals and mixed kaolinite-smectite minerals show decreasing prominence of the distinctive doublet absorption with increasing crystal disorder or decreasing kaolinite content. The polymorphs dickite and nacrite can be distinguished from kaolinite and halloysite based on doublet shape and band position. The vertical lines are placed at 1.4, 1.92, and 2.2 μm, and indicate absorptions in CRISM data that are diagnostic of kaolinite, halloysite or kaolinite-smectite clays. For details on spectra used, see Appendix. 122

Figure 8. CRISM spectra from K mica- (illite- or muscovite-) bearing materials. (a) unratioed CRISM spectra from FRT0000454E (thick line; 20.1°N, 70.1°E) and from 2x2 average spectra from HRS00002FC5 (thin line; 20.3°N, 70.2°E). (b) ratioed CRISM spectra using numerators and denominators from (a). Vertical lines at 1.4, 2.2, and 2.35 μm indicate absorptions due to overtones and combination tones of metal-OH related absorptions and likely indicate the presence of a K mica such as illite or muscovite. (c) Library spectra of K micas, illite and muscovite. For details on spectra used, see Appendix. 123

Figure 9. CRISM spectra from analcime-bearing materials. (a) unratioed CRISM ratio spectra from FRT00009312 (19.9°N, 65.9°E). (b) ratioed CRISM spectrum with numerators and denominators from (a). Vertical lines at 1.42, 1.79, 1.91, and 2.52 μm mark absorptions indicating the presence of analcime. (c) Library spectra of zeolites. For details on spectra used, see Appendix. 124

Figure 10. CRISM spectra from hydrated silica-bearing materials. (a) unratioed CRISM spectra from FRT00009312 (thick line; 19.9°N, 65.9°E) and from HRL00009EA3 (thin line; 16.8°N, 71.6°E). (b) ratioed CRISM spectra using numerators and denominators from (a). (c) Library spectra of basaltic glass with an opaline coating measured under Mars pressure and temperature conditions (reflectance scaled by a factor of 10) [Swayze et al., 2007] and opal-A/CT [Clark et al., 2007]. The vertical line at 1.91 μm shows the location of an absorption indicating the presence of H2O. Vertical lines at 1.39, 1.41, and 2.21 μm show the positions of absorptions that indicate Si-OH bonds and the presence of hydrated silica, e.g. opal, with differing degrees of hydration. For details on spectra used, see Appendix. 125

Figure 11. CRISM spectra from mixed hydrated silica-bearing materials. (a) unratioed CRISM ratio spectra from FRT0000A4FC (22.0°N, 77.2°E). (b) ratioed CRISM spectrum with the numerators and denominators from (a). (c) Library spectra of opal, zeolites, and polyhydrated sulfates for comparison [CRISM spectral library, measured by Cloutis et al., 2002; Clark et al., 2007]. Vertical lines are at 1.43, 1.93, 2.20, 2.26, 2.42 μm and, as discussed in the text, the observed absorption features do not unambiguously identify a single mineral but probably indicate a mixture of hydrated silica with either sulfates or zeolites. For details on spectra used, see Appendix. 126

Figure 12. CRISM spectra from magnesium carbonate-bearing materials. (a) unratioed CRISM spectra from FRT0000A09C (thick line; 21.3°N, 78.5°E) and FRT000040FF (thin line; 18.5°N, 77.4°E). (b) CRISM ratio spectra with the numerators and denominators from (a). (c) Library spectra of carbonates and hydrated silicates shown for comparison. The vertical lines in (a)-(c) are at

1.93, 2.31, 2.51 μm, respectively indicating the presence of H2O and the two absorptions due to carbonate. (d-f) same as (a-c) but over the 3-4 m wavelength range. The vertical lines in (d)-(f) are at 3.46 and 3.85 μm, indicating the two absorptions due to carbonate. The wavelengths impacted by a probable detector artifact are shown in gray. For details on spectra used, see Appendix. 127

Figure 13. Stratigraphy of kaolinite- and smectite-bearing units west of Nili Fossae. CRISM HRL0000B404 (21.8°N, 72.3°E) infrared albedo map colorized with parameter maps (R: BD1900H, G: BD2200, D2300) so that Fe/Mg smectite-bearing materials are magenta to blue and kaolinite-bearing materials are green. A closeup view of the contact between the cap, Fe/Mg smectite- bearing unit, and largest exposure with kaolinite is shown in a subset of HiRISE image PSP_008993_2020_RED. 128

Figure 14. Stratigraphy of kaolinite- and smectite-bearing units along Nili Fossae trough. (a) CRISM FRT00009971 (22.1°N, 74.6°E) infrared albedo map colorized with parameter maps (R: BD1900H, G: BD2200, D2300) so that Fe/Mg smectite- bearing materials are magenta to blue and kaolinite-bearing materials are green. (b) profile extracted from the MOLA DEM (CRISM DDR band 10) across the trough scarp from the line A-A’ in figure (a), showing a 900 meter scarp with a distinctive down-dropped block. (c) a subset of HiRISE image PSP_006989_2025_RED showing the black box in (a) at higher resolution. The kaolinite-bearing unit corresponds to the outlined brightest toned units which also have some blocks of low-calcium pyroxene (not shown). 129

Figure 15. Kaolinite-bearing units overlying sedimentary Fe/Mg smectite-bearing units. (a) MOLA digital elevation model with the outlines of CRISM image FRT0000A053 (22.4°N, 74.3°E) within the crater over the eroded central pit and FRT00009971 (Figure 14) over the Nili Fossae trough scarp to the southeast. The arrow indicates the location of the channel draining the crater (b) in MOLA elevation profiles, the crater of interest (green) is ~1 km shallower than nearby Hargraves crater (blue; 20.7°N, 75.8°E), a fresher crater of the same diameter, because of sedimentary fill. (c) CRISM FRT0000A053 infrared albedo map colorized with parameter maps (R: BD1900H, G: BD2200, D2300) so that Fe/Mg smectite-bearing materials are magenta to blue and kaolinite-bearing materials are green. The image was acquired over an area of eroded crater fill in the central pit, and the phyllosilicates are exposed by this erosion. (d) and (e) closeup views of exposures of kaolinite-bearing materials from the white boxes in (c) with HiRISE image PSP_007345_2025_RED. Materials which are kaolinite-bearing (green in (c)) are bright. (f) zoom in on (e) showing the lack of bedding in the bright kaolinite-bearing units. The arrow indicates fractures in overlying mantling material. 130

Figure 16. Stratigraphy of carbonate-bearing unit, kaolinite-bearing unit, and cap rock in eastern Nili Fossae. (a) CRISM FRT0000A09C (21.3°N, 78.5°E) false color composite image (R: 2.38, G: 1.80, B: 1.15 m) in which olivine is yellow, phyllosilicates are cyan, and carbonate is green. (b) Parameter map (R: BD2500, G: D2300, B: BD2200) where kaolinite-bearing units are blue to purple, Fe/Mg smectite-bearing units are green and Mg carbonate-bearing units are yellow to orange. (c) zoom view of the black box in (b). (d) subset of HiRISE PSP_002321_2015_RED showing the stratigraphy of carbonate- and kaolinite- bearing units beneath the cap rock. 131

Figure 17. Carbonate-, smectite-, and olivine-bearing rocks in eastern Nili Fossae. (a) False color CRISM images FRT00003E12, FRT0000B438, FRT0000A4FC, and FRT0000871C (R: 2.38, G: 1.80, B: 1.15 m) near 22.1°N, 77.1°E were used to colorize a portion of grayscale CTX image P03_002176_2024_XI_22N283W_070113. With this choice of bands, olivine appears yellow, Fe/Mg smectites are blue, and altered olivine (sometimes what is spectrally distinctly carbonate) is green. An arrow indicates a bright white-cyan patch which has a spectral signature of the mixed hydrated phase discussed in the text (silica+zeolite or silica+sulfate; Figure 11). A MOLA elevation map showing the topography of the scene is inset. (b) CRISM data from (a) was used to colorize a subset of HiRISE PSP_002888_2025_RED shown in the white box in (a). (c) same HiRISE subset colorized with a CRISM parameter map (R: OLINDEX, G:

BD2500, B: D2300). The red and blue channels are scaled by (Gmax-G/Gmax) so that carbonate-bearing units are green rather than white. Red is olivine and blue is Fe/Mg smectite. Non-carbonate altered olivine is magenta. (d) unratioed CRISM spectra from the units labeled 1-4 in panel (c). (e) zoom view of figure (c) showing the banded, partially altered olivine-bearing rocks beneath olivine sands. 132

Figure 18. Transported carbonate- and Fe/Mg smectite-bearing materials in Jezero crater. (a) False color CRISM image HRL000040FF (18.5°N, 77.4°E) with colors as in Figure 16. (b) Parameter map (R: D2300, G: BD2500, B: BD1900H). For each pixel, the red and blue channels are scaled by the value of the green channel ((Gscene max.-G)/Gscence max.) so that predominantly carbonate-bearing units appear green rather than white. Magenta to blue units are more Fe/Mg smectite-enriched. (c) CRISM ratioed spectra from regions of interest from light-toned deltaic sediments within the box shown in b. Spectra from each region were ratioed to the same denominator spectrum. 133

Figure 19. Mineral maps from a 25 km crater west of Nili Fossae near 20°N, 66°E. (a) CTX image mosaic colorized with mineral maps (R: BD2500, G: D2300, B: BD2200) from CRISM images FRT000051EE, FRT00009312, FRT00009E5D, FRT0000AC95, and FRT0000BD3A overlain. Materials bearing the zeolite analcime are red, hydrated silica are magenta, and Fe/Mg smectite and chlorite are green. The arrow shows the viewing perspective for (b). (b) three dimensional perspective view (3x vertical exaggeration) of the distribution of minerals within the crater provided by draping (a) over a MOLA digital elevation model. (c-e) subsets of HiRISE image PSP_003205_2000_RED showing in greater detail regions bearing hydrated silica, zeolite, and Fe/Mg smectite, respectively, taken from areas indicated in the inset images and in (b). (f) zoom of the area shown by the white box in (e) showing breccia blocks composing the central peak materials. 134

Figure 20. THEMIS day infrared mosaic with black arrows indicating the locations of TES/THEMIS identified granitoid/quartzofeldspathic materials from Bandfield [2006] along with minerals identified by CRISM. CRISM observation outlines and mineral detections within CRISM images are shown as in Figure 2. The dashed white box indicates the location of Figure 19. 135

Figure 21. Distribution of chlorite/prehnite- and K mica-bearing materials around a 50 km impact crater near 20°N, 69°E. (a) THEMIS daytime infrared mosaic with outlines of CRISM images. (b) CRISM mineral maps (R: BD2350, G: D2300, B: BD2200) in which K mica-bearing materials are magenta and chlorite/prehnite-bearing materials green to yellow (as explained in section 4.2, these minerals are difficult to distinguish). Yellow indicates a more prominent 2.35-μm absorption. Some Fe/Mg smectite may also be present in the scene and would also map as green; however, analysis of individual spectra show it is uncommon. (c) MOLA elevation profile across line A-A’ in (a). (d) subset of CTX image P04_002651_1999_XI_19N289W_070219 from the box in (b). (e) mineral indicators as in (b) overlain on (d) from CRISM images HRS00002FC5, FRT0000454E, FRT0000AACE, FRT0000A63F, and FRT0000B2C7. (f) chlorite/prehnite-bearing ejecta and polygonal ground from northeast of a small crater taken from the location shown by the white box in (e). 136

Figure 22. Craters in Northern Syrtis Major (near 16°N, 72°E). THEMIS day infrared mosaic is shown with parameter maps (R: D2300, G: BD2200, B: BD2350) from CRISM multispectral tile 1177 and FRT000050F2, FRT00009365, FRT00008D9A, HRL000086CA, FRT0000B1B5, HRL00009EA3, HRL0000B868, FRT00009786. White outlines show the coverage of the CRISM targeted (solid line) and mapping (dashed line) observations. Fe/Mg phyllosilicates present include include Fe/Mg smectites (red) and chlorite (magenta). Prehnite (magenta) is also present and does not map distinctly from chlorite in this color mapping. Around the southern crater, kaolinite (green) is found. Within the northern crater, hydrated silica (green) is found. The same color scheme is used for the targeted images to the right where the parameters are overlaid on a CRISM infrared albedo map. In the central peak of the northern crater in particular, there are mixed spectral phases (cyan, yellow) indicating chlorite, prehnite, hydrated silica, and possibly other hydrated phases spatially or intimately mixed in the same CRISM pixel. 137

Chapter 4: GEOLOGIC SETTING OF SERPENTINE DEPOSITS ON MARS

B.L. Ehlmann1, J.F. Mustard1, S.L. Murchie2

1Geological Sciences, Brown University

2Johns Hopkins University Applied Physics Laboratory

Published in

Geophysical Research Letters

March 25, 2010

137 138

Abstract

Serpentine, recently discovered on Mars using Mars Reconnaissance Orbiter data, is uncommon but found in three geologic settings: (1) in mélange terrains at the Claritas

Rise and the Nili Fossae, (2) associated with a few southern highlands impact craters, and

(3) associated with a regional olivine-rich stratigraphic unit near the Isidis basin. Any presently active serpentinization processes would be occurring beneath the surface and mineral products would not be apparent with surface and orbital data; however, finding serpentine in several Noachian terrains indicates active serpentinization processes in

Mars’ past. Important implications are the past production of magnetite, which may contribute to chemical remnant magnetization of Mars’ crust, and production of H2, which is a suitable energy source for chemosynthetic microbial life.

1. Introduction

Serpentines are 1:1 trioctahedral phyllosilicates with the general formula (Mg,

Fe)3Si2O5(OH)4 and form during hydrothermal alteration of ultramafic rocks at temperatures ranging from slightly above ambient to approximately 400°C. [e.g. Evans,

2004]. The presence of serpentine serves as a marker for distinctive aqueous chemical conditions: highly reducing, high pH, and low aSiO2. Serpentine’s existence on Mars is of significance for considerations of geophysics and habitability. Depending on fluid geochemistry during serpentinization, a commonly accompanying product is magnetite

(Fe3O4), which is highly susceptible to magnetism and has been hypothesized to contribute to the generation of Martian magnetic anomalies [Nazarova & Harrison, 2000;

Lillis et al., 2008]. Serpentinzation reactions also produce H2, which can serve as an 139

important energy source for chemosynthetic organisms [Schulte et al., 2006] or react abiotically with CO2 to produce methane [Oze and Sharma, 2005].

Data from orbiting visible/near-infrared imaging spectrometers have revealed phyllosilicate minerals in Mars’ oldest crust, the Noachian southern highlands. While most phyllosilicates found to date on Mars are Fe/Mg smectites or chlorites [Mustard et al., 2008, Poulet et al., 2005], which can form under a variety of conditions, less commonly occurring minerals such as recently identified serpentine [Ehlmann et al.,

2009] provide more specific information on past geochemical conditions. Here we catalogue the geologic settings of serpentine-bearing deposits found so far by the

Compact Reconnaissance Imaging Spectrometer for Mars (CRISM). We also discuss the prevalence of serpentine on Mars, likely constraints for its formation and persistence, and implications for geophysical and astrobiological studies.

2. Identification of serpentine-bearing rocks

Thousands of targeted, atmospherically corrected CRISM observations (18-

40m/pixel) [Murchie et al., 2009] have been examined in the southern highlands for evidence of alteration minerals, including serpentine. Serpentine is relatively uncommon.

To ensure no deposits had been overlooked in prior analyses, 162 observations in areas that show strong signatures of olivine [Poulet et al., 2007; Koeppen & Hamilton, 2008] or evidence for Fe/Mg phyllosilicates as indicated by the D2300 parameter [Murchie et al.,

2009] were re-examined in this study [for image processing methods, see Ehlmann et al.,

2009]. 140

Serpentine has characteristic absorptions in the 1.0-2.6 μm region that result from bending and stretching of metal(Mg,Fe,Al)-OH bonds [e.g., Calvin and King, 1997 and refs. therein] (Figure 1). The Mg-rich polytypes antigorite, lizardite, and chrysotile are indistinguishable at CRISM resolution. Their strongest absorption is an asymmetric, narrow 2.325 μm Mg-OH combination band. Near 1.4 μm, is a narrow OH stretch overtone, centered at 1.38 μm for Mg serpentines. A v-shaped band centered at 2.50-2.52

μm and a weaker feature at 2.10-2.12 μm, whose strength varies markedly with grain size

[Bishop et al., 2008], are also characteristic. The presence of all four of these absorption features is necessary to uniquely distinguish magnesian serpentines from other alteration minerals on Mars, such as carbonates, chlorites, Mg-rich smectites, and their mixtures

[Ehlmann et al., 2009]. Stoichiometric serpentine would not have molecular water in its structure, but spectra observed with CRISM, as well as many terrestrial serpentine samples, have a 1.9-μm H2O absorption caused by either intermixed hydrated phases or an imperfectly crystalline form. Iron-rich serpentines have a different set of absorptions

[Calvin & King, 1997]. Chamosite and berthierine have distinctive absorptions that were searched for and not found. Greenalite and cronstedite lack VNIR absorption features that uniquely distinguish them from other alteration products and their mixtures, and it is probably impossible to unambiguously determine their presence or absence on Mars from orbital data. On Figure 2, definitive and probable serpentine detections, all Mg-rich, are delineated. Probable serpentines have the characteristic shape and band positions of serpentine and are distinctive from other materials in the scene examined but may be lacking the weaker absorption features (Figure 1). 141

3. Geologic settings

All serpentine detections to date occur in rocks of three geologic settings: (1) in mélanges with other alteration minerals, (2) in the central peaks, walls, and ejecta of impact craters, and (3) in direct association with olivine-rich rocks.

3.1 Mélange terrains

Fourteen CRISM images were examined from the Claritas Rise, a faulted topographic high, which formed during Noachian, pre-Tharsis uplift and tectonism

[Anderson et al., 2001] and hosts a large crustal magnetic anomaly [Acuña et al., 1999].

Fe/Mg smectite and chlorite are the most common alteration minerals observed in nine images; three images show evidence of serpentine. In the largest exposures (26.8°S, -

101.2°E), raised knobs host serpentine-bearing rocks along with chlorite, kaolinite, and illite or muscovite (Figure 3a). Even at 5m/pixel resolution, no stratigraphic relationship or regular zoning of the minerals is observed, although the strongest signatures are found in local topographic highs and debris being shed from these.

West of the Nili Fossae (20.9°N, 73.3°E), where erosion has exposed rock beneath a dark-toned cap rock, CRISM spectra exhibit absorptions characteristic of serpentine, Fe/Mg smectite, kaolinite, low-calcium pyroxene, and olivine, each in discrete areas (Figure 3b). The mineral-bearing units represent a mélange of Noachian rocks with varying degrees of alteration. Examination of high-resolution imagery generally reveals no coherent stratigraphy or zoning. A possible exception is in a ~3-km2 area in the center of the CRISM image that shows a central area of serpentine ringed by

Fe/Mg smectite, although no distinct unit boundaries can be discerned. Throughout the scene, olivine is partially altered and exhibits 1.9- and 2.32-m absorptions which are 142

characteristic of either serpentine or saponite. This may indicate incipient alteration of olivine, or sub-pixel (<18m in scale) spatial mixing of olivine-bearing materials with altered materials.

3.2 Southern highlands craters

A small number of craters (~5% of those examined) scattered throughout the southern highlands exhibit spectral signatures consistent with serpentine, e.g. ejecta from a small crater within Chia crater, north of Valles Marineris (1.1°N, -59.8°E)(Figure 3c).

These craters are found in Noachian terrains By contrast, two-thirds of the craters in this study have materials bearing chlorite and Fe/Mg smectite. Serpentine is found in rocks of the central peaks, walls, and ejecta. Because of the small size of some of the exposures

(<5x5 CRISM pixels), it is often not possible to uniquely identify serpentine by the presence of characteristic 1.38- and 2.1-m absorptions. “Probable” serpentine is identified on Figure 1 where the spectral shape and band positions at 2.32 and 2.52 m differ from other materials in the crater and are consistent with the presence of serpentine.

3.3 Noachian bedrock in stratigraphic section

West of the Isidis basin, Early Hesperian Syrtis Major volcanic flows overlie

Noachian bedrock. The Noachian strata consist of an Fe/Mg smectite-bearing unit beneath an olivine-rich unit that in places has been partially altered to both carbonate and serpentine [Hamilton & Christensen, 2005; Mustard et al., 2009; Ehlmann et al., 2009].

North of the Noachian-Hesperian contact (17.3°N, 77.2°E), erosion exposes a >150-km2 outcrop of the olivine-rich unit, which is relatively bright-toned, highly fractured, and variably covered by dunes. The dunes, presumably derived from the bedrock, have 143

spectral signatures of olivine without spectral evidence for alteration products. However,

CRISM data indicate the olivine-rich bedrock composition here is heterogeneous, exhibiting partial alteration in most places to magnesium carbonate and in other places to serpentine (Figure 4). At the 5-m/pixel scale of CTX images, the morphology of the bedrock is similar regardless of alteration mineralogy. Examination of 25 cm/pixel

HiRISE images shows that it is fractured into meter-scale, angular blocks in a quasi- linear pattern with a fracture spacing of <2 m.

Elsewhere north of, in, and around the Nili Fossae, the olivine-rich bedrock also exhibits signs of partial alteration. In over 20 images, distinctive paired absorptions at

2.31 and 2.51m of comparable width and strength can be uniquely identified as magnesium carbonate [Ehlmann et al., 2008]. However, as noted by Mangold et al.

[2007] and Ehlmann et al. [2009], in places the olivine-rich unit instead exhibits characteristic absorptions at 1.91 and 2.32 m, due to H2O and Mg-OH respectively, and the 2.5-m absorption is weak to absent. This combination of absorptions is not consistent with carbonate but instead implies the presence of an Mg-rich phyllosilicate, either saponite or serpentine.

4. Discussion

Contiguous exposures of definitively serpentine-bearing materials are small and rare on Mars. This finding represents a combination of the actual distribution of serpentine and the fact that serpentine is difficult to distinguish uniquely from other

Fe/Mg phyllosilicates, carbonates, and their mixtures. To date, serpentine-bearing materials appear to be restricted to Noachian-aged rocks: mélange terrains near the 144

Claritas Rise and the Nili Fossae, a few impact craters throughout the southern highlands, and an olivine-bearing bedrock unit in Nili Fossae. That serpentine exists in these terrains confirms the existence of hydrothermal alteration of ultramafic rocks during the

Noachian epoch.

2+ Depending on temperature, fO2, and the activities of SiO2, H2O, Ca , serpentinization generates multiple additional products; brucite, talc, and magnetite are most typical [e.g. Frost & Beard, 2007]. Talc and brucite have distinctive IR absorptions but have not been observed, thus implying they are either absent or not optically dominant at IR wavelengths. Magnetite is largely opaque in the near-infrared and would not be detected by CRISM. Magnesium carbonates are a less common accompanying mineral in serpentinization reactions but can co-occur with serpentine in hydrothermal systems when the aCO2 is moderately high and temperatures are low. This may explain the occurrence of an olivine-carbonate-serpentine unit in the Nili Fossae region. Another possibility is a two-step process: olivine is partially altered to serpentine hydrothermally, and magnesium carbonate subsequently forms during surface weathering. On Earth, magnesite deposits are commonly observed in association with serpentinized ultramafic rocks as a result of surface or near-surface weathering [e.g. Keleman and Mater, 2008].

Olivine can weather directly to magnesite [Gooding, 1978; Jull 1988]; brucite is also highly susceptible to weathering in the presence of CO2 and is a common precursor mineral for hydrated and magnesium carbonates [Hostetler et al., 1966]. That the Nili-

Fossae olivine-carbonate-serpentine unit is heavily fractured on a decimeter to meter scale is consistent with rock fabrics in serpentinites, since the conversion from olivine to serpentine involves a 30-50% volumetric expansion [Evans, 2004 and refs therein]. 145

In Nili Fossae, a clear geologic record of process exists where both the reactants and products of serpentinization are found. The mechanism and timing of serpentine formation is less clear in the crater or mélange settings where rocks lack stratigraphic context. Serpentine may have formed prior to disruption, by cratering or whatever process created the mélange, or may have resulted from hydrothermal activity after the fact. Additional alteration minerals, albeit not directly associated with the serpentinization process, are sometimes found in nearby rocks and include kaolinite, illite and/or muscovite, chlorite, and Fe/Mg smectites. This suggests either mixing of pre- existing discrete units, perhaps by impact processes, or variation in precursor lithology or fluid composition and temperature on scales of just a few kilometers as in diatremes where kaolinite, illite, and chlorite form when pH drops and Eh increases making Al insoluble [Sader et al., 2007].

The global distribution of serpentine shows no correlation with magnetic field anomalies. Serpentine is found in Claritas and western Arabia, which have large magnetic anomalies; near the Nili Fossae, which has no anomaly; and not in Terra Sirenum nor

Terra Cimmeria where crustal magnetic sources are most intense. What controls the spatial distribution of the few serpentine occurrences on Mars? It is clearly not as simple as olivine distribution. Serpentine has not been definitively found in Argyre and Terra

Tyrrhena, which are regions of elevated olivine abundance [Koeppen & Hamilton, 2008], but multiple craters in apparently olivine-poor Arabia Terra have serpentine. Hydrated silicates are found in craters distributed in nearly all southern highlands terrains, suggesting that water availability was not a limiting factor for serpentine formation. Four plausible hypotheses are that (1) serpentine was once more abundant but has since mostly 146

weathered to other materials, e.g. Fe/Mg smectite; (2) serpentine distribution follows the distribution of buried, altered ultramafic rocks, which is different from surface composition measurable by orbiting sensors; (3) most olivine-bearing rocks are sufficiently Si-rich that the favored equilibrium hydrothermal alteration assemblage is chlorite- rather than serpentine-rich [Frost & Beard, 2007]; or (4) most alteration conditions were sufficiently oxidizing/low-T that other alteration products, e.g. oxides and smectites, were favored rather than serpentine.

5. Conclusions

Presently occurring serpentinization and consequent abiotic or biologically mediated methane production would be occurring in the subsurface and minerals produced could not be observed from orbit. However, it is clear that the serpentinization process operated on Mars in the past. In the Noachian period, formation of serpentine on

Mars near Nili Fossae, at the Claritas Rise, and associated with several southern highlands impact craters would have generated H2 and created aqueous habitats appropriate for chemoautotrophic microbial life. Evidence for such processes may be preserved at such sites. While most occurrences of serpentine are in disrupted strata, intact bedrock exposures such as near the Nili Fossae represent targets for future landed exploration. 147

Acknowledgements

Thanks to D. Cardace, W. Calvin, G. Swayze, S. Wiseman, M. Wolff and two anonymous reviewers for suggestions that improved this manuscript and to the MRO science operations teams for their efforts in collecting a spectacular dataset.

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151 Figures

Figure 1. CRISM spectra from the Claritas Rise (ratioed to nearby bland material to highlight the features) possess four distinctive near-infrared absorptions that permit the unique identification of serpentine[laboratory spectra from Clark et al., 2007]. An example spectrum of “probable” serpentine-bearing materials from a nearby scene has only some of the diagnostic features. 152

Figure 2. Global hillshade map of Mars from the Mars Orbiter Laser Altimeter (MOLA) indicating locations of serpentine-bearing materials and CRISM images analyzed in this study. 153

Figure 3. Serpentine in mélange and cratered terrains. CRISM parameter maps were used to colorize CTX images which show the morphology of materials bearing serpentine and other minerals (A) Claritas Rise, CRISM image FRT0000634B on CTX image P08_004108_1522_XI; (B) Nili Fossae, CRISM image FRT0000ABCB on CTX image P19_008347_2019_XN; (C) a small crater within Chia Crater, CRISM image HRL000095C7 on CTX image P15_006849_1815_XN. Colors of spectra correspond to colored areas on the images with those minerals. Spectra locations are indicated by arrows and circles. 154

Figure 4. Serpentine in stratigraphic section. (A) False-color (R: 2.38, G: 1.80, B: 1.15 μm) CRISM image HRL0000B072 on a CTX mosaic. The olivine-carbonate-serpentine unit (green) is being exposed from beneath Syrtis lavas (purple). (B) CRISM parameter map (R: OLINDEX, G: D2300, B: BD2500) where olivine appears red, carbonate appears white to pale yellow, and serpentine appears green. (C) the olivine unit and overlying lavas are shown in a subset of CTX image P06_003376_1987_XN. (D), (E) subsets of HiRISE image PSP_9217_1975 showing the meter-scale fracturing characteristic of the altered ultramafic unit. 155

Chapter 5: EVIDENCE FOR LOW-GRADE METAMORPHISM, DIAGENESIS, AND HYDROTHERMAL ALTERATION ON MARS FROM PHYLLOSILICATE MINERAL ASSEMBLAGES

Bethany L. Ehlmann1, John F. Mustard1, Roger N. Clark2, Gregg A. Swayze2, Scott L. Murchie3

1Department of Geological Sciences, Brown University, Providence, Rhode Island, 02912, USA

2U.S. Geological Survey, Denver, Colorado, 80225, USA

3Johns Hopkins University Applied Physics Laboratory, Laurel, Maryland, 20723, USA.

Submitted to

Clays & Clay Minerals

March 29. 2010

155 156

Abstract—The enhanced spatial and spectral resolution provided by the Compact

Reconnaissance Imaging Spectrometer for Mars (CRISM) on the Mars Reconnaissance

Orbiter (MRO) has led to the discovery of numerous hydrated silicate minerals on Mars, particularly in the ancient, cratered crust comprising the southern highlands. Recently discovered minerals include using visible/near-infrared spectra include: chlorite, prehnite, high-charge phyllosilicates (illite or muscovite), the sodium zeolite analcime, opaline silica, and serpentine. These minerals are not distributed homogeneously. Rather, certain craters host unusually diverse hydrated minerals which comprise distinctive alteration assemblages: (1) prehnite/chlorite-silica, (2) analcime-silica-Fe/Mg smectite-chlorite, and

(3) illite/muscovite-chlorite/prehnite. Furthermore, (4) serpentine has been found both associated with craters and in bedrock units. These assemblages contrast with prevalence of solely Fe/Mg smectites in most phyllosilicate-bearing terrains on Mars. Here, we review evidence that these mineral assemblages represent the result of aqueous alteration at elevated temperatures. We provide examples of minerals’ geologic occurrence from west of the Isidis basin near Nili Fossae, and discuss differences in implied temperature, water-rock ratio, fluid composition, and starting materials during alteration. Of the minerals found to date, prehnite provides the clearest evidence for subsurface, hydrothermal alteration, since it forms only under highly restricted conditions (e.g., T =

200-350ºC). Multiple mechanisms (some at elevated temperature, others not) exist for forming the other individual minerals. However, for the characteristic assemblages observed, we suggest that from basaltic starting materials, the probable formation mechanisms are, for (3), transformation of trioctahedral smectites to chlorite and dioctahedral smectites to illite during diagenesis; for (1) and (2), low temperature 157

(<400ºC) hydrothermal circulation of fluids in basalt; and for (4) (<400ºC) hydrothermal circulation of fluids in ultramafic rocks. Evidence for high-grade metamorphism at elevated pressures or temperatures >400ºC has not been found.

Key Words—Mars; infrared spectroscopy; phyllosilicates; zeolites; hydrothermal alteration; metamorphism; diagenesis; analcime; prehnite; serpentine; illite; muscovite; chlorite; silica 158

Introduction

High spatial and spectral resolution orbital visible/near-infrared spectroscopy of

Mars has greatly expanded capabilities for detection of alteration mineral phases and examination of their geologic context. The Observatoire pour la Minéralogie, l’Eau, les

Glaces, et l’Activité (OMEGA) on Mars Express has acquired near global coverage at hundreds of meters per pixel resolution and revealed deposits of phyllosilicates, which are mostly smectite clays, as well as sulfates (Bibring et al., 2005; Gendrin et al., 2005;

Poulet et al., 2005). Data from the Compact Reconnaissance Imaging Spectrometer for

Mars (CRISM) onboard the Mars Reconnaissance Orbiter (MRO) have subsequently revealed additional diversity in alteration mineral phases including carbonates, zeolites, additional sulfates, and additional hydrated silicates, such as kaolinite, chlorite, serpentine, and opaline silica (Mustard et al., 2008; Milliken et al., 2008; Ehlmann et al.,

2008; 2009).

The presence of these alteration minerals is alone interesting as an indicator of past liquid water chemically reacting with Mars’ basaltic rocks and sediments. In some instances, it is possible, however, to gain further information about the precise geochemical conditions of aqueous alteration (temperature, pressure, pH, Eh, and ion activities) by considering the nature of the minerals, their assemblages, and their geologic context in conjunction with knowledge about the thermodynamic and kinetic limitations on mineral formation, derived from terrestrial experience. OMEGA found that phyllosilicates were restricted to Mars’ oldest Noachian terrains (>3.7 Gyr) while sulfates were found in younger Hesperian terrains (approximately 2.9-3.7 Gyr). This suggested a 159

directional evolution to Mars’ history whereby initially neutral to alkaline conditions in which there was long-term water rock interaction needed to form Fe/Mg smectites were replaced by a more ephermerally wet, more acid environment where precipitation of sulfate evaporites dominated (Bibring et al., 2006). CRISM findings have largely supported this paradigm but have also permitted the discrimination of at least 11 distinct alteration environments on the basis of morphology and mineral assemblage (Murchie et al., 2009a). Here we focus on a subset of alteration minerals recently identified by

CRISM, which suggest that in Mars’ past, some aqueous alteration took place at elevated temperatures. We review evidence that localized hydrothermal, low-grade metamorphic, or diagenetic activity were some of the processes forming and modifying hydrated silicate minerals on Noachian Mars.

Water-rock reactions at elevated temperature require available heat and water.

Ancient Mars (>3.5 Gyr) possessed both surface (Fassett & Head, 2008) and subsurface

(Andrews-Hanna et al., 2007) waters. Sources of heat were available from impact cratering, volcanism, and a planetary geothermal gradient somewhat higher than that today. Hence, numerous authors have predicted the hydrothermal or metamorphic alteration of the Martian crust by water-rock interactions. Evidence includes trace quantities of primary hydrous minerals, amphiboles and micas, in the SNC meteorites, carbonate isotopic values in ALH84001, and enriched D/H ratios in the Martian crust and atmosphere (Gooding et al., 1992; Harvey & McSween, 1996; Jakosky & Jones, 1994).

The occurrence and mineralogy of hydrothermal systems depend on water availability, the composition of the host rocks, and the temperature and composition of circulating fluids. Based on experience with terrestrial systems, the existence of phyllosilicates, 160

zeolites, carbonates and sulfates have been predicted in hydrothermal systems generated in Mars’ basaltic rocks following impacts (Newsom, 1980; Allen et al., 1982). Models have shown that, given sufficient water in the regolith, impact-induced hydrothermal systems could last up to tens of thousands of years in the subsurface beneath craters

(Rathburn and Squyres, 2002; Abramov and Kring, 2005; Barnhart et al., 2008) producing zoned phyllosilicate-bearing mineral assemblages (Schwenzer and Kring,

2009). Hydrothermal activity has also been invoked to explain some apparently water- related features around volcanoes such as incised valleys on Martian stratovolacano slopes (Gulick & Baker, 1990; Farmer, 1996 and references therein). The expected mineralogic assemblages resulting from hydrothermal circulation beneath lava flows have been calculated (Griffith & Shock, 1997). Some crustal cooling models indicate hydrothermal activity in the upper 10 km of crust may have played a substantial role in heat loss on early Mars (Parmentier and Zuber, 2007). Finally, one possible explanation for the existence of methane in the present day Martian atmosphere is release as a product of ongoing serpentinization, the reaction of ultramafic rocks and water at elevated temperatures (Atreya et al., 2007).

Below, we begin by describing the methodology for delineating indicator minerals for metamorphism, hydrothermal activity, and diagenesis and then detecting them remotely from Mars orbit. We then provide examples of the geologic setting of four diagnostic minerals/mineral assemblages found on Mars that we believe are indicative of alteration at elevated temperatures: (1) prehnite/chlorite-silica, (2) analcime-silica- chlorite-smectite, (3) chlorite and illite/muscovite, and (4) serpentine. We describe what each of these may indicate in terms of past temperature, pressure and geochemical 161

conditions and make some suggestions for future research and the potential of these types of sites as targets for future landed exploration.

Methods

Delineating Diagnostic Minerals

A metamorphic facies is “a set of metamorphic mineral assemblages…showing a regular relationship between mineral composition and bulk chemical composition, such that different metamorphic facies…appear to be related to different metamorphic conditions, in particular temperature (T) and pressure (P), although other variables such as PH2O may also be important” (Smulikowski et al., 2003). Figure 1 shows the classic distribution of these facies by temperature and pressure. Mars is a basaltic planet, so for each facies, mineral products from basaltic protoliths (Table 1) are most relevant to the study of metamorphism on Mars. For any given facies, minerals of metamorphosed terrains may be divided into common, diagnostic, and forbidden (Chen et al., 1983 after

Zwart et al., 1967). Common minerals are those typically found within a facies.

Diagnostic minerals specifically indicate particular P, T values during alteration via thermodynamic restrictions on the conditions in which they form. Assemblages of diagnostic+common minerals provide routine field and petrologic means of identifying metamorphic facies and thus particular P,T conditions experienced by the rock.

Forbidden minerals are those which definitively do not belong in a particular metamorphic facies, most likely because they definitively indicate particular P,T conditions of another facies. 162

For relatively low temperature alteration, which as discussed below is most relevant to the mineral phases discovered on Mars to date, there are fuzzy boundaries between sub-greenschist facies metamorphism, zeolite facies metamorphism, hydrothermalism, and diagenesis with different authors adopting different conventions through time (see Arkai et al., 2003 for a review). Sub-greenschist metamorphism occurs at temperatures less than approximately 300-400°C and pressures approximately typically

<4 kbar but extending to 8 kbar (Schiffman & Day, 1999). The facies is defined by the presence and stabilities of pumpellyite and prehnite.

The lowest metamorphic grade is the zeolite facies. Zeolites are attributed to alteration at low temperatures in the presence of CO2-poor or -absent aqueous fluids.

Zeolites bridge the gap between diagenetic and metamorphic assemblages. Some authors have divided the facies into diagenetic, including heulandite or analcime (plus quartz), and metamorphic, including laumonite. However, after Coombs et al. (1959), the zeolite facies is now regarded as a single metamorphic facies regardless of the origin of the mineral (Arkai et al., 2003). The upper temperature is approximately 300°C and the precise assemblage depends on P, aSiO2, and PH2O (Coombs et al., 1959). Commonly, zeolite facies are only locally developed. Complete chemical equilibrium is not achieved for the entire rock because of the low temperatures of metamorphism. Furthermore, unmetamorphosed basalt is nearly anhydrous, and H2O must be added to the rock to produce low-grade metamorphic assemblages. The rock must be sufficiently permeable to fluid entry and, therefore, zeolites are typically only found in fractures, vugs, and other cavities. Numerous types of zeolites are found and have been used in conjunction with 163

thermodynamic data to determine temperatures of alteration in terrestrial environments

(e.g. Weisenberger & Selbekk, 2008).

Finally, diagenesis refers primarily to thermal- and pressure-related maturation of sediments, rather than that of rocks. In particular, smectites come into equilibrium with their surroundings, reordering into other aqueous minerals with time and temperature when water is available to promote transformations. Dioctahedral smectites such as montmorillonite convert to mixed layer illite-smectite clays, then to illite, then to white micas such as muscovite with more advanced diagenesis at higher temperatures.

Trioctahedral smectites such as saponite, with increasing grade, transform to corrensite (a

1:1 mixed-layer chlorite/smectite) then to chlorite (Merriman & Peacor, 1999 and refs. therein).

Using infrared spectroscopy to identify diagnostic minerals/assemblages

Visible/near-infrared spectroscopy over the 0.4-2.5 m wavelength range permits the detection and differentiation of numerous minerals (e.g. Clark et al., 1990) and has been used extensively in the characterization of phyllosilcates, carbonates, and zeolites

(e.g. Bishop et al., 2002a, 2002b). Broad electronic absorptions, resulting from charge transfer and ligand field transitions, indicate the presence of transition metals, most commonly iron, in the mineral structure while sharper combination and overtone vibrational absorptions indicate the presence of H2O, metal-OH, and CO3 in the mineral structure.

The shape and position of several diagnostic bands are used to identify a given mineral. The presence of H2O in the mineral structure is expressed near 1.9 μm by a combination tone of the fundamental bending and stretching vibrations of the water 164

molecule. An overtone of the metal-OH stretching vibration occurs near 1.4 μm, as does an H2O combination tone. From 2.1-2.5 μm, additional structural metal-OH combination stretching plus bending vibrations are found, and their precise wavelength position depends on the type of octahedral cation and mineral structure. For example, Fe-OH absorptions in smectites are found near 2.28 μm whereas Al-OH absorptions are found near 2.20 μm. The position of the OH overtone near 1.4 μm also shifts depending on composition (e.g. Clark et al., 1990; Bishop et al., 2002a; 2002b).

Overtones/combinations of fundamental vibrations related to H2O in zeolites (Cloutis et al., 2002) and CO3 in carbonates (Gaffey, 1987) are also diagnostic and found in the 2-

2.6 μm wavelength region.

Not all minerals produce diagnostic spectral signatures, however. For example, of the minerals in Table 1, quartz and feldspars do not possess identifiable features in the

VNIR (an exception is when these phases are hydrated, as in opaline silica or contain substantial iron impurities as in lunar Fe-bearing plagioclases). Furthermore, typically, in natural rocks and sediments, only 1-2 phases may be apparent in a single spectrum even though the area sampled may be comprised of far more than two minerals. When mixed areally (“checkerboard mixing”) minerals contribute linearly to the spectrum; however, in semi-transparent or fine-grained samples, “intimate mixing” dominates instead. This is normally the case in remote sensing applications and hence minerals comprising surface rocks and sediments contribute non-linearly to observed spectral properties, out of proportion to their abundance (Clark, 1999). Nevertheless, by considering the VNIR- detected minerals in given areas and nearby areas on the surface, the mineralogic 165

assemblage characteristic of a particular lithologic unit or family of units can sometimes be determined.

Fortunately, several metamorphic minerals diagnostic of particular facies have distinctive near-infrared spectroscopic signatures (Figure 2). Some zeolite facies minerals can be clearly distinguished from other phases by relatively deep water-related absorption features (e.g. at 1.16, 1.42, 1.79, 1.91, 2.13, and 2.52 m in analcime). Prehnite has distinctive sharp-OH features, in particular from 2.2-2.4 m and at 1.48 m. The amphiboles indicative of higher degrees of metamorphism have similar spectra to one another but can be clearly differentiated from the minerals diagnostic of lower metamorphic grades. These characteristics can be used in remote sensing of Mars to identify metamorphic grades from orbital data, and similarly by examination of other minerals, hydrothermal alteration and diagenesis.

Identifying and Mapping Minerals on Mars

CRISM is a hyperspectral imaging spectrometer on the Mars Reconnaissance

Orbiter (MRO) with 544 channels that sample the visible to near-infrared spectral region from 0.4 - 4.0 μm (Murchie et al., 2007). CRISM operates in either a 72-channel mapping mode that will provide global coverage at 200 m/pixel or a full 544-channel targeted mode that provides 10 x 10-20 km images at a resolution of 15-38 m/pixel.

Spectral data from 1.0-2.6 μm in the latter, high-resolution mode are the focus of this paper and allow identification and mapping of mafic minerals and alteration minerals of highest interest.

Prior to spectral analysis, atmospheric and photometric corrections were applied to CRISM data cubes to correct for viewing geometry and separate out the contribution of 166

reflected light from the surface from the effects of atmospheric transmission (Ehlmann et al., 2009; Murchie et al., 2009b). Following atmospheric correction, a noise removal algorithm that removes data spikes in both the spectral and spatial domains but does not affect broader absorptions of mineralogic interest was implemented (Parente, 2008). In addition, to highlight spectral differences between areas and reduce residual atmospheric and instrumental artifacts in spectra, average spectra from regions of interest were divided by spectra from a dusty or spectrally unremarkable region in the same scene.

Once a particular mineral or group of minerals is identified, their spectral properties may be parameterized and mapped spatially, e.g. by calculating absorption band depth for each pixel in an image or by more complex feature-fitting algorithms

(Pelkey et al., 2007; Clark et al., 2003). These mineral mapping techniques have been employed with success in terrestrial remote sensing, including in metamorphic terrains

(e.g. Kozak et al., 2004; Brown et al., 2006; Swayze et al., 2009). The CRISM parameter maps were coregistered with and combined with a variety of imaging datasets to obtain information about the geomorphology and geologic context of mineral detections. These include coordinated observations acquired by the two high-resolution MRO cameras, the

Context Imager (CTX) and the High Resolution Science Experiment (HiRISE) cameras on MRO, which acquire images at 5 m/pixel and 0.25 m/pixel, respectively (Malin et al.,

2007; McEwen et al., 2007). Collectively, information on diagnostic minerals and orbital data from Mars were used to identify numerous areas where evidence of likely low-grade metamorphic, hydrothermal or diagenetic activity is preserved.

167

Observations

Mineralogy of the Martian Surface

Table 2 summarizes the hydrated silicate alteration phases found on Mars to-date.

These alteration minerals are restricted to the southern highlands and Noachian terrains.

Most are exposed by cratering, associated with thousands of craters across the southern

highlands, although a few eroded escarpments (Nili Fossae, Mawrth Vallis, Valles

Marineris, and Uzboi Vallis) provide excellent exposures of intact stratigraphic cross-

sections. Regional surveys of cratered terrains in Terra Tyrrhena, Noachis Terra,

(Fraeman et al., 2009) and west of the Isidis basin in cratered terrains and in northern

Syrtis Major and near the Nili Fossae (Ehlmann et al., 2009) were conducted to examine

the prevalence and distribution of alteration mineral phases. These studies have shown

that the most common alteration minerals are Fe/Mg smectites (Figure 3; Mustard et al.,

2008). In a few instances, these smectites are accompanied by additional alteration phases. Chlorite-bearing or chlorite/prehnite-bearing materials are the most common of these. Rarer are zeolites, hydrated silica, and muscovite/illite. Serpentine has also been rarely found and mapped on Mars (Ehlmann et al., 2010).

Fe/Mg smectites can form in a wide range of terrestrial environments from pedogenic (e.g. weathering of basalt), to hydrothermal, to lacustrine precipitation. Their detection on Mars indicates past aqueous alteration under neutral to alkaline pH, but does not permit elucidation of the details of the conditions under which alteration took place

(temperature, pressure, further details on fluid composition). However, as discussed above, other minerals can be more diagnostic of the pressure, temperature, and geochemical conditions at the time of formation. The locally diverse alteration minerals 168

likely indicate locally distinctive aqueous processes. In particular, observations of four diagnostic minerals (or mineral assemblages) provide diagnostic evidence for alteration at elevated temperatures (Table 2).

Prehnite-bearing rocks

Prehnite occurs in multiple locations in ancient Martian cratered terrains (Clark et al., 2008), including the greater Argyre basin region (Buczkowski et al., 2008), the craters at the northern margin of Syrtis Major (Ehlmann et al., 2009), and in Noachis

Terra and Terra Tyrhenna (Fraeman et al., 2009). Prehnite forms in hydrothermal or metamorphic conditions and has distinctive OH-related absorptions centered at 2.35-2.36

μm with additional weaker bands at 2.24 and 2.28 μm (Figure 4). Prehnite’s principal

VNIR absorption coincides with the location of a 2.33-2.35 μm absorption found in chlorite, which makes it challenging to discriminate them remotely. However, prehnite has an OH overtone located at 1.48 μm, apparently unique among all other minerals

(Clark et al., 2007). Chlorite has a different, distinctive overall spectral shape and an OH overtone at 1.39 μm. The presence of the 1.48 μm absorption permits the unique discrimination of prehnite’s presence. Interestingly, where found on Mars to date, prehnite is always in a mixture with chlorite.

Figure 4 is a typical exposure of prehnite-bearing rock. Within a relatively fresh

40 km crater (near 16°N, 72°E), prehnite/chlorite is found along with Fe/Mg smectites and silica (Figure 4; Ehlmann et al., 2009). Nearly all of the rocks in and around the central peak of the crater are altered and there are numerous mixed pixels which show attributes of multiple alteration minerals, i.e. are mixtures of phases. Materials in the central peak are heavily brecciated and are excavated from depth (<4km for a crater of 169

this size, based on models by Melosh, 1989) by the impact process. Prehnite/chlorite and

Fe/Mg smectite are the areally most widespread phases. Spectra of endmembers are shown in Figure 4. Silica (discussed further in the next section) is found in smaller areas in rocky knobs. All alteration minerals are found both in the crater central peak as well as in the wall rock and crater ejecta.

Analcime-Silica-Smecite-Chlorite assemblages

Analcime is rarely found on Mars but a few craters north of the Hellas basin and in northern Syrtis Major near the Antoniadi Basin have collections of diverse alteration minerals which include analcime. Phases also commonly found near analcime-bearing materials include silica and Fe/Mg smectite (Ehlmann et al., 2009). Spectra of zeolite minerals are dominated by water-related absorptions. Analcime has strong absorptions at

1.42, 1.91, and 2.52 μm and weaker absorptions at 1.16, 1.79, 2.14, and 2.52 μm.

Analcime forms in highly alkaline environments, either in saline-alkaline lakes or in low water:rock ration hydrothermal alteration and metamorphism.

Silica has been found in diverse geologic settings on Mars from layered deposits in and around Valles Marineris to southern highlands craters (Milliken et al., 2008;

Ehlmann et al., 2009). Opaline silica and hydrated basaltic-glass share similar H2O and

Si-OH absorptions near 1.4, 1.9, 2.2 μm. In opaline silica, phases with more hydrogen bonded H2O have broader bands at these positions as is the case for silica-bearing materials in Figure 5 (vs. Figure 4; see Swayze et al., 2007, Milliken et al., 2008, for measurements). Silica forms at a variety of pHs and temperatures and is not by itself diagnostic of a formation environment (McLennan, 2003) 170

A 25 km impact crater (20°N, 66°E) near the Antoniadi basin hosts diverse mineral deposits in and around the central peak (Ehlmann et al., 2009). Rocks comprising the central peak are brecciated and both the blocks and the matrix are altered. Depending on the precise area of the central peak considered, rocks have as the spectrally dominant alteration mineral analcime, Fe/Mg smectite, and chlorite (Figure 5). Ringing the central peak are rippled sands of debris reworked by aeolian processes. These deposits host most of the hydrated silica. Additional spectral signatures of silica (and smectite) are found in a few isolated knobs of rock on the crater wall. This is one of two craters in the region with the same mineral assemblage.

Chlorite/Illite bearing materials

Chlorite is the second-most common class of phyllosilicate on Mars, after Fe/Mg smectite. It is often found mixed at a sub-pixel scale (<18m/pixel) with prehnite-bearing materials (e.g. Figure 4 above) or in terrains associated with Fe/Mg smectites. As discussed above, the spectral signature of chlorite is distinguished by its unique overall shape as well as absorptions at 2.33-2.35 m, 1.39 m, and a shoulder at 2.26m (King and Clark, 1989). In pure terrestrial specimens, the 1.9 m is absent. Chlorite can be a primary mineral in granitoid rocks and also forms via hydrothermal, metamorphic, and diagenetic reactions.

The spectral signatures of the high interlayer charge phyllosilicates illite and muscovite have very similar features at 1.40, 2.20, 2.235, and 2.44 m (Kruse and Hauff,

1991) that cannot be distinguished at CRISM spectral resolution (Ehlmann et al., 2009).

Illite/muscovite is not commonly observed on Mars and is found in only a very few impact craters (Fraeman et al., 2009). Muscovite can be magmatic, hydrothermal, or 171

metamorphic. Illite can form in hydrothermal or diagenetic reactions or by weathering of muscovite.

Chlorite- and illite/muscovite-bearing materials are usually spatially associated on

Mars with deposits of illite/muscovite accompanied by nearby chlorite (Ehlmann et al.,

2009; 2010). Unlike mixing of chlorite and prehnite, which is typically at a sub-pixel scale, chlorite- and illite/muscovite-bearing materials can be separated more readily into spatially discrete exposures. In Figures 6 and 7, we observe in the walls and ejecta of a 50 km crater (near 20°N, 69°E) the distinctive spectral signatures of these two minerals. The chlorite-bearing materials apparently have some additional contribution from prehnite- bearing materials from the presence of the 1.48 m absorption (Figure 6). The presence of chlorite is definitive, given the spectral shape. The areally dominant materials in this crater’s ejecta are chlorite- and prehnite-bearing. Illite/muscovite occurs in small knobs comprising a few CRISM pixels (Figure 7), yet still the spectral signature of illite/muscovite can be isolated with careful analysis (Figure 6). Many pixels exist which are mixtures of chlorite and illite/muscovite in variable proportions (Figure 6, middle spectrum). The 2.35 m absorption increases in strength as the chlorite component increases. No pixels are found with just a 2.2 m absorption, as might be the case were the mineral responsible another Al-phyllosilicate such as montomorillonite or kaolinite.

As the 2.2 m band increases in strength, the 2.35 m band shows a correlated increase, also indicating the single illite/muscovite phase is the most probable explanation for the data.

The strongest absorptions are coincident with raised knobs and ridges (arrows in

Figure 7). The signatures become diffuse as mass wasting and eolian erosion transport 172

and comminute the hydrated mineral-bearing materials and mix them with local dusts and soils. These deposits and their knobby appearance may result from fluid flow in conduits which become mineralized and then more resistant to subsequent erosion. Alternatively, these are materials in the crater’s ejecta blanket and may be rocks that have been excavated from depths <5km by this impact and then differentially eroded.

Serpentine-bearing rocks

Serpentine has been identified on the surface of Mars in the ejecta and central peaks of craters, in terrains that are a chaotic mixture of units with multiple alteration minerals but no obvious stratigraphy, and in an olivine-rich stratigraphic unit south of the

Nili Fossae. A distinctive spectral shape and four distinctive absorptions at 1.38, 2.11,

2.325, and 2.52 m spectrally discriminate serpentine from other Martian alteration minerals (Figure 8e; Ehlmann et al., 2010). Serpentine results from the hydrothermal alteration of ultramafic rocks at temperatures ranging from slightly above ambient to approximately 400°C (e.g. Evans, 2004). Fluids circulating in the rocks are high pH, low aSiO2, and are typically highly reducing (Frost & Beard, 2007).

Figure 8 exhibits the location of serpentine where it is found in stratigraphic section, associated with an olivine-rich unit (17°N, 77°E). This unit, which has been mapped west and south of the 1900 km Isidis Basin, has a banded, fractured appearance, drapes underlying Fe/Mg smectite-clay rocks, and is estimated to have >~30 wt.% olivine

(Hoefen et al., 2003; Hamilton & Christensen, 2005; Mustard et al., 2007; 2009). Here the olivine-rich unit is exposed beneath younger Hesperian lavas of the Syrtis Major formation. The olivine-rich unit is partially altered, in some locations exhibiting spectral signatures indicative of the presence of serpentine and in others exhibiting the spectral 173

signature of magnesium carbonate, which is probably a weathering product from either the olivine or the serpentine (Figure 8; Ehlmann et al., 2009; 2010). To date, no other sites on Mars have been found where serpentine is part of a coherent statigraphy (further details are provided in Ehlmann et al., 2010).

Discussion

Hydrothermal alteration/low-grade metamorphism of mafic and ultramafic rocks

Among the diverse alteration minerals mapped on Mars, prehnite provides the clearest evidence for aqueous alteration at elevated temperatures. Figure 9 shows stability diagrams for prehnite in the context of other metamorphic facies and minerals. Prehnite formation is restricted to a relatively narrow range of conditions. Thermodynamics and experimental data show it is a stable to metastable product over 200-400°C and at <3kbar

(a slightly greater range than shown on Figure 9a; Robinson & Bevins, 1999; Schiffman

& Day, 1999). In addition to temperature and pressure constraints, during prehnite formation, the mole fraction CO2 in the circulating fluids must be low, otherwise the Ca would be incorporated into calcite and prehnite would not form. Calcite has thus far not been definitively identified in the southern highlands of Mars, although it has been found in the soils at high latitudes (Boynton et al., 2010).

That there appear to be no “pure” prehnite VNIR spectral signatures at 18m/pixel scale in the areas studied and that it is always identified with chlorite is also explained by examination of predicted metamorphic assemblages. At relatively low pressures and temperatures less than <~400°C, prehnite is expected to co-occur with chlorite (Figure 174

9a). Indeed, the prehnite/chlorite+silica assemblage observed in the crater in Nili Fossae is an expected low-grade, low-pCO2 metamorphic assemblage (Figure 9b).

At lower grades but with similar protoliths, analcime occurs as one of the minerals in zeolite facies metamorphism (Figure 9a), and it is often accompanied by silica. Unlike prehnite, there is no lower bound to analcime formation, so its detection alone does not indicate aqueous alteration must have occurred at elevated temperatures.

In alkaline lake sediments where water interacts with detrital clay and volcanic glass, analcime can precipitate from ~250°C down to ambient temperatures. In lacustrine environments, the common mineral assemblage is clinoptilolite, erionite, analcime, and sodium silicates and borosilicates (e.g. Eugster et al., 1980). In the analcime-bearing craters observed on Mars (e.g. Figure 5), these other hydrated silicates have not been found nor is there evidence for lacustrine sedimentary units disrupted by cratering.

Rather, instead silica and smectites are the more common accompanying alteration minerals. The morphologic and mineralogic evidence collected to date is collectively more indicative of diagenetic or hydrothermal settings for analcime on Mars. In Icelandic hydrothermal systems, for example, analcime is one of a suite of zeolites that serves as a

“geothermometer”, indicating temperatures of hydrothermal alteration between 75-225°C in altered basalts (Wiesenberger and Selbekk, 2008). It is commonly accompanied by

Fe/Mg smectites, chlorites, silica, and other zeolites.

We suggest that the VNIR spectral signatures and morphologic setting of the prehnite/chlorite-silica and analcime-silica-smecite-chlorite assemblages may best be explained by hydrothermal alteration of basaltic rocks. The temperatures and pressures implied are at the boundary of what is variously termed low-grade metamorphism and 175

hydrothermal alteration of basalt, depending on the fluid source and water:rock ratio.

Regardless of the terminology used, the mineralogy implies alteration at elevated temperatures, above 200°C but probably below 400°C. The pressures probably do not exceed 3 kbar. Where serpentine is found instead, the pressure and temperature conditions implied are the same, although the ultramafic rather than mafic precursor rock results in different fluid composition precipitation of different minerals (Ehlmann et al.,

2010). Lower aSiO2 and aCaO means serpentine is favored over chlorite and prehnite.

Magnesium carbonate may form in the same hydrothermal system as serpentine when fluids are in communication with the CO2-rich atmosphere or, alternatively, Mg carbonate may form as a later weathering product (Ehlmann et al., 2008; 2010).

No amphiboles indicative of higher-pressure (blueschist) or higher temperature

(greenschist) metamorphic facies have been found, even though these phases are infrared active and would be detected by CRISM if present at sufficient abundance in the rock at an 18m/pixel scale in the areas surveyed (Figure 2). The apparent absence of phases such as actinolite and glaucophane suggests that higher-grade metamorphism has not been an important process in Mars history, a conclusion that makes sense given the absence of the typical tectonic settings in which these minerals form on Earth. The transient high heats and pressures of impact may be insufficiently long-lived for these minerals to form

(Schwenzer & Kring, 2009). Furthermore, given Mars lower gravity, much greater depths are required to reach the pressures for blueschist facies amphibole formation, and these materials are rarely excavated by cratering (Figure 1).

Figure 10 shows terrestrial examples of the nature of rocks that could produce the spectral signatures observed in and around the southern highlands craters discussed here. 176

In these basaltic rocks, fluid flow through vugs and fractures has resulted in mineralization of zeolites and prehnite-pumpellyite while the bulk rock itself is only weakly altered (usually to Fe/Mg smectite and mixed layer chlorite-smectite). When

VNIR spectral signatures of altered basalts are ratioed to unaltered basalts, the VNIR spectral signatures of the alteration minerals will be prominent, even though they are in vugs and may comprise only a small portion of the surface area. Differential weathering of the precipitated phases may serve to concentrate these depending on their relative hardness. For example, when subjected to aeolian erosion and transport, precipitated silica nodules will be resistant to breakdown. Silica will tend to remain behind as a lag while clays and zeolites are broken down into finer-grained particles which can be removed by transport. As previously noted by Ehlmann et al. (2009), this may explain the occurrence of silica in aeolian deposits ringing the central peak of the 25 km crater in

Figure 5 and may also explain the occurrence of zeolites in the dust of Mars (Ruff, 2004).

Alternatively, the silica might be explained by the aqueous weathering of glassy impactites.

Diagenesis

On Earth, smectites rarely persist in the long-term rock record because diagenetic processes transform these clays to illite or chlorite. That smectite clays are the predominant clay on Mars has been used to argue that the integrated contact time of these rocks and sediments with fluids must be short (Tosca & Knoll, 2009). So far, the detection of illite/muscovite is rare on Mars and restricted to disrupted strata. In the case of the 50 km crater in Figure 7, it is apparent that the mineral signatures are closely associated with knobby topographic highs. However, it is difficult to tell whether the 177

illite/muscovite- and chlorite/prehnite-bearing rocks constitute the ejecta of the crater or whether the entire eastern side of the crater has been heavily eroded and the chlorite/prehnite and illite/muscovite simply represent exposed bedrock (the deposits shown are in a broad topographic low some 800m topographically below the regional level). In either case, rocks from depth in the crust are probably being exposed.

In the broader region of Mars around this 50 km crater, Fe/Mg smectites are the dominant alteration minerals, found in breccias and in sedimentary deposits (Ehlmann et al., 2009; Figure 3). Kaolin group minerals or kaolin-smectite mixed layers are also found as a thin unit capping the smectites (Ehlmann et al., 2009). The chlorite- and illite/muscovite-bearing materials found east of this 50 km crater (Figure 7) and at other craters may have formed by the diagenesis of these Fe/Mg- and Al-clays at depth. On

Earth, muscovite, can be primary in granitic terrains and weather to illite. However, on

Mars where basalt is dominant, illite/muscovite more likely results from hydrothermal activity or diagenesis of Al-rich clays. For dioctahedral smectites, diagenesis begins upon burial, usually as temperatures reach 50-80ºC and illite/muscovite phases become more ordered with temperature and time. For trioctahedral smectites like saponite, the reaction proceeds instead through chlorite-smectite mixed layers to chlorite (Merriman and

Peacor, 1999). Given the predominance of Fe/Mg smectites on Mars, chlorite would be expected to dominate a diagenetic assemblage. Yet the proportion of chlorite on Mars due to diagenesis rather than hydrothermal activity will remain difficult to assess without landed investigations.

178

Outstanding questions

As CRISM continues to map the surface of Mars, additional exposures of diverse mineral assemblages may yet be found that provide further insights on the conditions of aqueous alteration. Certainly, heat and water were available on ancient Mars to generate the low-temperature and low-pressure hydrothermal, metamorphic, and possibly diagenetic conditions in many geographic areas, including those reviewed here. A key advance in interpretation of VNIR spectra would be the ability to discern remotely the presence of mixed layer clays such as illite-smectite and smectite-chlorite that indicate lower temperature alteration and diagenesis. Some progress is being made in this regard

(Milliken et al., 2010).

Cooling of the early Mars crust (Parmentier and Zuber, 2007), heat from impacts

(Newsom, 1980), or heat from volcanic flows, e.g. Syrtis Major, may have supplied the heat necessary for the elevated temperature alteration observed. In the case of crustal cooling, different minerals would have precipitated stratified by depth, with the phases governed by the geothermal gradient and pressure, e.g. zeolites shallow, prehnite deeper, amphiboles deeper still. Interestingly, the diagnostic minerals mapped so far appear to show no characteristic stratification with depth or differences in materials excavated as a function of crater diameter (larger craters excavate materials originally deeper in the crust) (Fraeman et al., 2009). This suggests either that crustal cooling was not the mechanism responsible or that repeated churning by impacts into the crust has destroyed the evidence for mineral stratification. The mineral assemblages found also do not appear to have any strong geographic regional dependencies (Figure 3). For example, the craters of figures 4-6 are all within ~500km of each other yet exhibit strikingly different mineral 179

assemblages. Controls on the spatial distribution of alteration minerals are not well- understood at this time.

Exposures of Fe/Mg smectites in intercrater sedimentary deposits and topographic lows in Terra Sirenum lack any evidence of associated minerals formed at elevated temperatures. Most of the exposures of diagnostic minerals and mineral assemblages are in craters rather than intact stratigraphies. This makes further interpretation of the geologic setting of alteration difficult. When associated with craters, a key question is the timing of mineral formation. Did the low-grade metamorphic/hydrothermal processes pre-date or post-date formation of the impact craters? Did shock metamorphism play a role? Modeling shows that post-impact hydrothermal activity is restricted to the subsurface beneath crater (interior, including the central peak) and only if a crater lake is sustained do waters reach the surface (Rathburn and Squyres, 2002; Abramov and Kring,

2004). Hence, minerals in crater walls and ejecta presumably formed prior to the cratering event, while the timing for those in the central peak could be more ambiguous.

Some authors predict shock temperatures near central peaks would be too high for hydrated silicates to survive; therefore, minerals found (e.g. Figure 4, 5) within craters must be hydrothermal in origin (Fairen et al., 2009). Other authors model the ballistic ejection of materials without alteration and also model lower temperatures and pressures experienced in the central peaks (Barnhart et al., 2008; 2010). Resolving the origin of the low-grade metamorphic/hydrothermal assemblages requires merging high-resolution models of impact cratering processes (e.g. Abramov and Kring, 2005) with geochemical models (thermodynamic and kinetic) for mineral formation (e.g. Schwenzer and Kring,

2009) and destruction. By coupling detailed knowledge of temperature and mass 180

movement with understanding of the rates of mineral formation/destruction, the question of pre- or post-impact alteration can be rigorously assessed.

This question is important for understanding the timing and duration of aqueous activity on early Mars. Was hydrothermal activity restricted to the Noachian and the time of impact-basin formation and initial crustal cooling? In this instance, alteration minerals excavated by craters are exposed at later times, but hydrothermal systems were only active very early in the planet’s history (>~3.9 Ga, the age of the last major basin). The alterative is hydrothermal alteration, perhaps related to small impacts or perhaps to emplacement of volcanic flows extending from the Noachian into the Hesperian period of

Mars (>2.7 Ga). Such longevity would bode well for the sustenance of habitatable environments through time. Terrestrial hydrothermal systems host diverse microbial communities, including some of the most deeply rooted (ancient) organisms on the phylogenic tree. Intact, coherent stratigraphies of Martian hydrothermal alteration (e.g.

Figure 8), represent prime targets for future landed missions searching for biosignatures or mineralized organic matter.

Summary and Implications

Some of the hydrated silicate minerals detected on Mars by CRISM serve as geothermometers for the temperature of aqueous alteration and indicate hydrothermal and/or low-grade metamorphic conditions. Prehnite provides definitive evidence for alteration at higher temperatures (200-400°C) and typically co-occurs in association with chlorite and silica. Analcime commonly occurs in assemblages with silica and Fe/Mg smectite. In these assemblages, analcime too likely indicates hydrothermal alteration, 181

although at somewhat lower temperatures (50-200°C) than prehnite. Serpentine indicates similar temperature conditions of alteration (<400°C) but of ultramafic rather than mafic precursor rock. Finally, the occurrence of illite/muscovite in association with chlorite- bearing materials may be the product of diagenesis (<100°C) of smectite-bearing sediments. In all cases, the alteration conditions implicated are low-grade. High temperature and pressure phases indicating higher grades of metamorphism have not been found on Mars. Key outstanding questions include the nature of the processes that generated the hydrothermal environment (e.g. source of heat) and their timing. Further studies of the nature, timing, and longevity of Martian hydrothermal systems will enable understanding how the availability of liquid water and planetary habitability changed through time.

Note: CRISM images may be previewed and downloaded via a map search tool at http://crism.jhuapl.edu. The image IDs of all spectra shown are provided in short form in the figures with the leading zeros of the eight-digit image ID omitted, e.g. image 454E is

0000454E.

182

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Tables

Table 1. Common diagenetic mineral assemblages and diagnostic and common minerals for select metamorphic facies for a basaltic protolith (see Spear, 1995, Ch. 11 for an overview and also specific references below).

P-T Assemblage/ Diagnostic Common Minerals and Assemblages Specific References Facies Minerals and Assemblages Diagenesis Chlorite, illite/muscovite, and mixed layer clays (chlorite- Cann, 1979; Meunier, smectite, illite-smectite); celadonite+Fe/Mg smectite 2005

Zeolite zeolite heulandite + analcime + quartz, Arkai et al., 2003; minerals laumontite, stilbite, wairakite Coombs, 1959 Subgreenschist prehnite, prehnite+pumpellyite+chlorite, Frey & Robinson, 1999 (prehnite- pumpellyite prehnite+chlorite+quartz pumpellyite) epidote+calcite+quartz+chlorite Greenschist actinolite, chlorite+albite+actinolite(cummingtonite) Smulikowski et al., 2007; cummingtonite +epidote+ zoisite+quartz Philpotts & Ague, 2009 Amphibolite hornblende hornblende+albite+epidote+chlorite, Smulikowski et al., 2007; horneblende+Ca Philpotts & Ague, 2009 plagioclase(+quartz+garnet) Blueschist glaucophane, glaucophane+lawsonite+garnet+quartz, Smulikowski et al., 2007; lawsonite glaucophane+lawsonite, Philpotts & Ague, 2009 glaucophane+lawsonite+jadeite

193

Table 2. Hydrated silicate minerals resulting from aqueous alteration, discovered on Mars through January 2010. All mineral phases are observed from orbit with VNIR spectroscopy. Those which serve as diagnostic indicator minerals for elevated temperatures and diagenesis, hydrothermal alteration, or low-grade metamorphism are indicated in the leftmost column.

Indicator, Hydrated silicate Formula References elevated T group/mineral Fe/Mg smectites Poulet et al., 2005; (e.g., nontronite, (Ca, Na) (Fe,Mg, Al) (Si, Al) O (OH) 0.3-0.5 2-3 4 10 2 Mustard et al., 2008 saponite) Poulet et al., 2005; Montmorillonite (Na,Ca) (Al,Mg) (Si O )(OH) 0.33 2 4 10 2 Mustard et al., 2008 Kaolin group minerals (e.g. Bishop et al., 2008; Al Si O (OH) kaolinite, 2 2 5 4 Ehlmann et al., 2009 halloysite) Mustard et al.,., 2008; Chlorite (Mg,Fe2+) Al(Si Al)O (OH) 5 3 10 8 Ehlmann et al., 2009

Serpentine (Mg,Fe)3Si2O5(OH)4 Ehlmann et al., 2009 High charge Al,K phyllosilicate Mustard et al., 2008; KAl AlSi O (OH) (e.g. muscovite or 2 3 10 2 Ehlmann et al., 2009 illite) Clark et al., 2008; Prehnite Ca Al(AlSi O )(OH) 2 3 10 2 Ehlmann et al.,., 2009

Analcime NaAlSi2O6·H2O Ehlmann et al., 2009 Milliken et al.,, 2008; Opaline silica SiO ·H O 2 2 Ehlmann et al., 2009

194

Figures

Figure 1. Pressure-temperature metamorphic facies diagram after Spear (1995). 195

Figure 2. Example visible near-infrared spectra of diagnostic metamorphic minerals. These were acquired in a laboratory (Clark et al., 2007). Distinctive shapes and positions of absorption features permit these minerals to be identified on the surface of Mars using CRISM spectral data in the 1.0-2.6 μm wavelength range. 196

Figure 3. A Mars Orbiting Laser Altimeter (MOLA) hillshade map of the topography of Mars with mineral detections from a survey of CRISM targeted images overlain. Symbols represent minerals found within a single CRISM image at a particular location (this spatial scale is too large to display the images). Fraeman et al. (2009) surveyed Noachis Terra and Terra Tyrrhena. Ehlmann et al., (2009) surveyed the area north of Syrtis Major. Fe/Mg smectites are the most commonly detected alteration mineral, although there are localized instances of higher mineralogic diversity. The area outlined by the white box is the focus of this review. 197

Figure 4. A 40 km crater in northern Syrtis major (16°N, 72°E) with prehnite/chlorite- and silica-bearing rocks. These minerals along with Fe/Mg smectite can be mapped in the crater walls, central peak, and ejecta. CRISM mineral maps were constructed by mapping absorption band depths at 2.3 m (red), 2.2 m (green) and 2.35 m (blue) on a Context Image composite mosaic. CRISM spectral data in comparison to data from laboratory samples (Clark et al., 2007; Swayze et al., 2007) were used to verify the mineral maps. Some pixels are composed of prehnite mixed with chlorite. Others have the distinctive signature of a hydrated silica phase. 198

Figure 5. Analcime, silica, Fe/Mg smectite, and chlorite have been identified in a 25km in northern Syrtis Major (20°N, 66°E). CRISM mineral maps were constructed from the band depth at 2.5 m (red), 2.3 m (green) and 2.2 m (blue) and verified by comparing CRISM spectra to data from spectral libraries (Clark et al., 2007). The mineral maps are overlain on a mosaic from the Context Imager and projected using gridded MOLA topographic data at 3x vertical exaggeration. 199

Figure 6. Spectral signatures of CRISM chlorite-prehnite and illite/muscovite and their mixtures compared to laboratory data from Clark et al. (2007). All spectra were taken from the images mapped in Figure 7. 200

Figure 7. Chlorite/prehnite and illite/muscovite minerals found in the walls and ejecta of a 50 km impact crater in northern Syrtis Major (20°N, 69°E). CRISM mineral maps were constructed from images 2FC5, 454E, and A8C6 by mapping absorption band depths at 2.35 m (red), 2.3m (green), and 2.2 m (blue) and are shown overlain on a CTX image. The strongest mineral signatures coincide with rocky knobs and become weaker with distance from these knobs. Spectra from these CRISM images are shown in Figure 6 and were used to verify mineral identifications indicated by the color mapping. 201

Figure 8. Serpentine in stratigraphic section in northern Syrtis Major (17°N, 77°E). Serpentine is found associated with an olivine-rich unit that is variably altered to Mg carbonate and serpentine. (a) A subset of CRISM image B8C2 in false color (red: 2.38, G: 1.80, B:1.15 um) overalin on a Context Imager image shows the olivine-rich unit (yellow-green) beneath younger lavas (dark purple unit on the left side of image). (b) a THEMIS daytime infrared mosaic shows context: this CRISM image covers part of the eastern boundary of the lava flows. (c-e) ratioed CRISM spectra taken from the locations indicated by the arrows in (a) allow discrimination of olivine, Mg carbonate, and serpentine within the rock unit. 202

Figure 9. Phase diagrams indicated prehnite stability as a functions of (a) temperature and pressure. Metamorphic facies (ZEO=zeolite, GS=greenschist, BS=blueschist) are indicated and select reactions are shown (adapted from Schiffman & Day, 1999). (b) Temperature and mole fraction CO2 (adapted from Robinson & Bevins, 1999). Prh=prehnite, Pmp=pumpellyite, Chl=chlorite, Anl=analcime, Qtz=quartz, Alb=albite, Lws=lawsonite, Zo=zoisite, Tr=tremolite, Ep=epidote, Act=actinolite, Cal=calcite. 203

Figure 10. Basaltic rocks altered by hydrothermal activity/low-grade metamorphism from (a) zeolitization, western Iceland and (b) and prehnite/pumpellyite and silica precipitation from hydrothermal alteration of basalts of the southwestern Newfoundland Bay of Islands ophiolite. 204

Chapter 6:

WEATHERING AND HYDROTHERMAL ALTERATION

OF BASALTS IN ICELAND: MINERALOGY FROM VNIR,

TIR, XRD, AND IMPLICATIONS FOR LINKING MARS

ORBITAL AND SURFACE DATASETS

B. L. Ehlmann1, J. F. Mustard1, and D. L. Bish2,

1Dept. of Geological Sciences, Brown University, 2Dept. of Geological Sciences, Indiana University

204 205

ABSTRACT

There is growing evidence for both pedogenesis and hydrothermal activity on Mars at neutral to alkaline pH, indicated by the detection from orbit of mineral assemblages of smectites, opaline silica, and zeolites from orbit using visible/near-infrared (VNIR) spectroscopy. Here we examine basaltic rocks from Iceland, collected from basalt flows that were in some places altered at the surface and in other locations were hydrothermally altered by non-sulfurous groundwater circulation as a geochemical analog for Noachian

Mars. Rock samples were surveyed in the field using a portable VNIR spectrometer.

Altered and unaltered rocks that were typical for the locality were collected as were altered rocks whose spectra were most similar to those measured by CRISM from Mars orbit. Samples were selected for detailed laboratory analyses included black to reddish and friable to coherent basaltic rocks and basaltic rocks with fractures various filled by zeolites, silica, and celadonite. Analyses included: (1) measurement of reflectance spectra of the whole rock by the ASD; (2) measurement of reflectance spectra of particle-size separates derived from the bulk rock and from minerals extracted from the vesicles; (3) measurement of X-ray diffraction (XRD) patterns that were analyzed to determine quantitative modal mineralogy; and (4) electron microprobe chemical analyses. These data emulate orbital data from infrared spectrometers (e.g. TES, CRISM, OMEGA), which detect the infrared active components, linked to in-situ data such as will be measured by the ChemMin XRD/XRF instrument on the MSL rover. While XRD was found to provide the most complete determination of sample modal mineralogy, VNIR spectroscopy was the technique most useful for determination of the composition of 206

smectite minerals present, demonstrating the key synergies between these two techniques relevant to the study of aqueous alteration on Mars.

Introduction

Recent orbital investigations have revealed distinctive assemblages of minerals that indicate aqueous alteration on early Mars took place in diverse environments

(Mustard et al., 2008; Murchie et al., 2009), including low-temperature/pressure hydrothermal systems at neutral to alkaline pH (Ehlmann et al., 2009a). In recent investigations of terrestrial Mars-analog sites, neutral to alkaline pH alteration of basalt has been neglected in favor of sulfur-rich, acidic systems including those at the Hawaiian volcanoes and Rio Tinto, Spain, which may be typical of Hesperian (3.1-3.7 Gyr)

Martian environments (e.g. Morris et al., 2005; Fernandez-Remolar, et al., 2005;

Hurowitz & McLennan, 2007).

In cratered Noachian terrains on Mars, mineralogic assemblages of Fe/Mg smectites, chlorite, opaline silica, and zeolites such as analcime have been detected

(Ehlmann et al., 2009a). Sometimes one or more of these phases is accompanied by prehnite, a material that definitely indicates alteration at T>200°C (Frey and Robinson,

1999; Ehlmann et al., 2009). Elsewhere, aluminum clay-bearing rocks have been found stratigraphically above iron-magnesium clay-bearing rocks (Wray et al., 2008; Bishop et al., 2008; Ehlmann et al., 2009; Loizeau et al., 2010). These two settings have been inferred to represent evidence of hydrothermal alteration and pedogenesis (top-down leaching) on early Mars, respectively. 207

In terrestrial geological investigations of paleoenvironment, a number of techniques are typically employed, including complete characterization of rock modal mineralogy, examination of small-scale textures, and bulk geochemical analyses.

Geologic investigations of Mars’ paleoenvironments take place in a comparatively data- poor environment. Hydrated silicate minerals were first detected on Mars using the

Observatoire pour la Minéralogie l’Eau les Glaces et l’Activité (OMEGA), an orbiting visible/near-infrared (VNIR; 0.4-5 m) spectrometer (e.g. Bibring et al., 2005; Poulet et al., 2005). The conclusion that early Mars likely hosted hydrothermal and pedogenic environments is based largely on qualitative mineral identifications at 18 m/pixel scale, obtained from the Compact Reconnaissance Imaging Spectrometer for Mars (CRISM), an second orbiting VNIR (0.4-4 m) imaging spectrometer (Murchie et al., 2007). The orbiting Thermal Emission Spectrometer (TES; Christensen et al., 2001) has the capability of modeling whole rock mineralogy, but perhaps because of the larger footprint (3 km x 5km) or a difference in sensitivity at thermal wavelengths (7-50 m) to alteration minerals, the alteration phases apparent to CRISM are not apparent to TES.

The basaltic rocks of Iceland provide an ideal setting and samples with which to examine how conclusions draw from infrared spectroscopy of whole rocks (such as acquired from Mars orbit) compare to the more detailed information that can be obtained via in-situ techniques such as can be employed in terrestrial laboratories and on landed missions to Mars. We began study of the alteration of basalt lava flows in Iceland as a geochemical analog for Noachian (> 3.7 Gyr) Mars. The basaltic bedrock in Iceland is recently formed (<16Myr) and <10% by volume is of more highly evolved compositions.

The land surface has poorly formed soils and sparse vegetation, and most Icelandic 208

hydrologic systems are fed by waters that are meteoric, rather than derived from seawater. Hence, ground and surface waters may be broadly similar to those that may have existed on Noachian Mars. Iceland has a variety of geothermal spring systems (low

T, low S; low T, high S; and high T, high S) each of which has distinctive pH, Eh, and aqueous geochemistry and creates distinctive mineralogic assemblages (Franzson et al.,

2008). Here we examine rocks collected from basalt flows that were in some places altered at the surface by pedogenesis and in other locations were hydrothermally altered by non-sulfurous groundwater circulation (low T, low S). Our sample characterization methods were to compare mineralogy determined with infrared spectroscopy

(comparable to Mars orbital data) to information derived from in-situ data (chemistry,

XRD) such as will be measured by the 2011 Mars Science Laboratory rover.

Understanding the linkages between remote spectroscopic and in-situ data will be critical to successful traverse planning, data interpretation, and furthering understanding of the paleoenvironments of Mars.

Geologic setting

Rocks from two localities were studied here (Figure 1). The Hvalfjordur basalt succession consists of 3-4 Myr plateau basalts, subject to zeolitic alteration (Neuhoff et al., 1999; Franzson et al., 2008; Wiesenberger & Selbekk, 2008). The Berufjordur plateau sequence is older, probably ~ 9-11 Myr, based on age dating of other members of the plateau sequence at a nearby site (Kissel et al., 2010) and exhibits similar zeolitization.

The two study areas differ from younger (<3.3 Myr) volcanic successions principally by their lack of hyaloclastite formations, which are characterized by the presence of 209

palagonite and probably formed during volcano-ice interactions after the onset of glaciation (Franzson et al., 2008). These have previously been examined as Mars analogs

(Bishop et al., 2002a). The samples in this study are all volcanic rocks which have been buried by accumulation of younger volcanics. For these samples, alteration occurred at relatively low temperatures at the surface; in heated environments near the surface, driven by emplacement of hot overlying lavas from subsequent eruptions; and/or by burial and heating by the geothermal gradient. Both study sites are not near active or relict volcanic sources, and thus altering fluids were not sulfurous. Glacial and fluvial erosion provide access to the deeper parts of the basalt plateau sequence.

Methods

The general approach was to select samples on the basis of their VNIR spectral characteristics, with particular focus on those which were similar to CRISM orbital VNIR spectra acquired from Mars orbit. The samples were then carefully analyzed in the laboratory. Reflectance spectra were acquired of particle size separates from the bulk rock and fracture/vesicle fill at VNIR and TIR wavelengths and used to qualitatively determine mineralogy. XRD and flux fusion elemental analysis were used to quantitatively determine mineralogy and elemental chemistry.

Infrared spectroscopy (VNIR, TIR)

Rock samples were surveyed in the field using an Analytical Spectral Devices

(ASD) portable VNIR spectrometer over the 0.4-2.5 m wavelength range. Altered and unaltered rocks that were typical for the locality were collected as well as altered rocks whose spectra were most similar to those measured by CRISM from Mars orbit. ASD 210

whole rock spectra were then reacquired in the laboratory with a contact probe attachment. Spectra were acquired from multiple portions of each rock then averaged.

Samples from the bulk rocks as well as samples of minerals extracted from vesicles and fractures were ground, dried, and sieved into particle-size separates (<25 m

(A), <150 m (C), and 75- or 106-500 m (D)). Reflectance spectra were measured over

0.4-25.0 m for the particle size separates in the Keck/NASA Reflectance Experiment

Laboratory (RELAB; Pieters, 1983), using both a bidirectional spectrometer (measured relative to Halon and corrected for absolute reflectance) and a Nicolet 740 Fourier- transform spectrometer (measured relative to a rough gold surface in a H2O- and CO2- purged environment and corrected to reflectance).

Over the 0.4-4.0 m VNIR wavelength range, spectra were examined for broad electronic absorptions, which indicate the presence of transition metals, most commonly iron, and sharper combination and overtone vibrational absorptions, which indicate the presence of H2O, metal-OH, and CO3 in the mineral structure (e.g. Burns, 1993; Clark et al., 1990). Over the 5.5-25.0 m TIR wavelength range, emissivity was approximated from the reflectance data by Kirchoff’s law (E=1-R), an approach which has been shown to yield good agreement with actual emissivity measurements (Mustard & Hays, 1997;

Bishop et al., 2008b). This was done to best emulate data from the orbiting TES and rover-based mini-TES instruments. Spectra were examined for fundamental vibrational absorptions due to Si-O, metal-Si-O, and H2O and OH structural elements. The measurement of multiple particle size separates permitted evaluation of any grain size effects on spectral properties of these samples.

211

XRD

X-ray diffraction (XRD) was performed on the <25m size fraction samples over a 2 range of 2º to 70º at 0.02º 2 intervals using a Bruker D8 Advance diffractometer with Cu K radiation, incident- and diffracted-beam Soller slits, and a solid-state PSi detector. All samples were initially measured with count times of 2s/0.02º2 step, and selected samples were re-run with 15s count times to improve counting statistics.

Additionally, all samples were measured with 15s count times over the 2 range of 58º to

64º to measure the d(060) peak. The b-axis spacing allows distinguishing dioctahedral montmorillonite from trioctahedral saponite and dioctahedal nontronite (Moore &

Reynolds, 1997).

To verify the presence of smectite and check for possible interstratified or non- expanding clays, untreated samples with a diffraction peak near 15Å were prepared as oriented mounts, dried, saturated with ethylene glycol, and then measured from 2-20º 2.

Incorporation of ethylene glycol into the interlayer of smectites causes the d(001) spacing to expand to ~16.7Å, whereas non-expanding clays such as chlorite, illite, and kaolinite do not experience a peak shift. Patterns from oriented, glycolated samples were also examined for irrationality of higher-order reflections that would indicate the presence of interstratification.

Areas of the most intense peaks from component minerals were measured, and the relative percentage of each constituent was determined using the reference intensity ratio

(RIR) method (Chung, 1974), using literature RIR values. Rietveld refinement was also applied, using the Topas Rietveld program, to determine the relative abundance of the well-ordered (i.e., non-clay minerals) components. The Rietveld method allows 212

refinement of unit-cell parameters for component minerals and it also provides information on crystallite size and strain contributions to peak broadening. Contributions from smectite were not modeled as the mineral is not three-dimensionally ordered and is not properly modeled by the Topas program. Final refined abundances represent weight percents obtained from the Rietveld method scaled to include the RIR-determined abundance.

Elemental Analysis

Flux fusion of powdered samples followed by ICP-AOES was used to measure the weight percent of Al, Ca, Fe, K, Mg, Mn, Na, Si, and Ti using the procedures outlined in Murray et al. (2000). For each sample (<150 m size fraction), 40mg of ground and dried sample was mixed with lithium metaborate flux in a 4:1 flux to sample ratio in a graphite crucible. The mixture was then fused in a furnace at 1050°C for 10 minutes. The fused bead was poured into a bottle with 10% nitric acid and agitated for > 30 minutes to promote dissolution. The sample solution was then filtered through a 0.45 micron filter to remove graphite and diluted to 4000x for analysis by ICP-AOES. Unknown samples were run in triplicate along with 3 blanks and 8 standard reference samples. ICP data were corrected for drift over the measurement run and for matrix effects using the blanks, and then were converted to elemental weight percent. These data were then multiplied by a correction factor to account for losses due to the flux fusion technique. The factor was derived by linear regression of standard weight percent element by ICP measured weight percent for the eight standard references. 213

Results

Results from both VNIR spectroscopy and XRD indicate that the samples can be divided into two categories: (1) bulk rock samples in which smectite is the dominant alteration mineral, and zeolites, SiO2 phases, and other phyllosilicates are present in only minor amounts; and (2) samples extracted from veins and vesicles in which alteration minerals other than smectite are the dominant alteration mineral and alteration minerals comprise most of the sample.

VNIR

VNIR spectra of all samples show evidence for aqueous alteration, including the

1.9m combination tone from the stretch and bend of the H2O molecule (Figures 3-5).

The strength of the ~1.9-μm band in spectral data is related to the degree of hydration, both in water molecules in the interlayer and surface hydration; band strength can vary depending on the humidity, thermal history of the smectite, and the nature of the water

(Milliken and Mustard, 2005). Additional intermixed hydrated phases, e.g. zeolite, may also enhance the 1.9-μm band without introducing additional bands in the wavelength region from 2.0-2.6 μm. OH and metal-OH overtones and combination tones are also visible near 1.4m and over the 2.0-2.6m region. Some minerals composing the sample, usually two or less, can be identified using the VNIR data (Table 1).

Whole rock spectra

All whole rock spectra acquired with the ASD field spectrometer show evidence for alteration as indicated by absorptions at 1.4 and 1.9 m. There is also substantial variation in the character of the Fe-related electronic transition absorptions at wavelengths <1.5 m. Some prominent vibrational absorptions may be seen longward of 214

2 m, including at 2.2 m in Hvalfj017 and Wtrfall016 and at 2.3 m in Icel009/010.

Other spectra have more subtle features that are challenging to identify confidently above the level of noise in the ASD’s longest wavelength measurements. Hence portions of all rocks were prepared for further VNIR spectral measurements in the laboratory.

Bulk rock samples

Five samples of bulk rock have spectra consistent with a mixture of high-calcium pyroxene and Fe/Mg smectite (Figure 4). Pyroxene exhibits broad electronic transition absorptions, centered near 1.0 m and 2.2 m for high-calcium pyroxenes like augite

(Burns, 1993). The 1.0 m band is prominent in these samples while the 2.0 m is weaker and appears to become less prominent with a contribution from alteration minerals.

In smectite clays, the relative proportions of Fe, Mg, and Al cations in the octahedral site dictate the presence and position of bands from 1.35-1.45 μm, 2.2-2.3 μm, and near 2.8 μm. The 2.3-μm band results from combination tones of the 2Fe-OH and

3Mg-OH bends and stretches, and its position varies depending on the relative proportions of Fe vs. Mg in the Fe/Mg smectite solid solution. For example, nontronite, the dioctahedral Fe endmember, has a band center near 2.29 μm (Grauby et al., 1994;

Bishop et al., 2002c; Frost et al., 2002), whereas the trioctahedral Mg endmembers saponite and hectorite have absorptions near 2.31-2.32 μm (Figure 4b) (Clark et al.,

1990). Aluminum, if present in smectite, is indicated by a strong 2.21 μm 2Al-OH combination tone (Bishop et al., 2002b, 2008b).Combination tones of H2O are present at

1.41 μm, as is a structural OH stretching overtone that shifts in position, depending on the octahedral cation (Bishop et al., 1994): at 1.43 μm for 2Fe-OH (Bishop et al., 215

2002a,b; Frost et al., 2002), at 1.41 for 2Al-OH (Bishop et al., 2002a,b), and at 1.38-1.39

μm for 3Mg-OH (Clark et al., 1990; Bishop et al., 2002a, 2002b). The fundamental of the metal OH stretch lies at 2.70 μm for Mg-OH, 2.76 μm for Al-OH, and 2.80 μm for Fe-

OH (Figure 6b).

The band positions of these OH related features in the bulk rock samples are given in Table 2. The continuum removed band depth near 2.3 m ranges from just greater than zero (Hvalfj025) to 2% (Icel009, Icel010). Based on the position of the position of the 2.3-m feature, all samples are intermediate Fe/Mg smectites between nontronite and saponite (Figure 6a). Hvalfj025 and Hvalfj054 may be more iron-rich based on the band positions at 1.4 and 2.3 m. From the 2.3 m, Icel009 and 010 appear to be more Mg-rich but also have a strong 2.8-m absorption due to Fe-OH. Hvalfj011 may be a mixture of multiple smectites; two distinct features due to Fe-OH and Mg-OH are observed near 2.3 m while the Al-OH band is most prominent near 2.76 m.

The final bulk rock sample, Hval017 appears to be a mixture of hematite and the

Al smectite montmorillonite, with an Mg or Fe-enriched phase specifically indicated by an Al-OH band centered somewhat longward of 2.2 m at 2.215 m (8% strength;

Bishop et al., 1994) and a 2.44 m band. Hematite is indicated by an electronic absorption at 0.86 μm and a charge transfer absorption at 0.53 μm.

Vesicle and fracture fill

From the four samples from rock veins and vesicles, only single minerals can be identified from VNIR data (Table 1). Hvalfj023 appears to contain the dioctahedral mica celadonite, with characteristic absorptions of approximately equal strength at 2.30 μm

(Fe3+Fe3+OH combination) and 2.35 μm (MgMgMgOH combination) and strong slope to 216

1.0 μm (Figure 5). A weaker absorption can also be seen at 2.26 μm (AlFe3+OH combination) (Bishop et al., 2008b).

NIR spectral features of zeolites are dominated by sorbed water (Hunt and

Salisbury, 1973; Clark et al., 1990), and are the likely source of broad and deep 1.4 and

1.9 μm features in Hvalfj055 and Hvalfj057 (Figure 5). Sometimes, due to their different structures and cations, zeolites have distinctive absorptions that permit discrimination of distinct minerals (Clark et al., 1990; Cloutis et al., 2002). Hvalfj057 has analcime, which is characterized by strong absorptions at 1.42 μm, 1.91 μm, 2.52 μm and progressively weaker absorptions at 1.79 μm, 1.12 μm, 2.14 μm. Sample Hvalfj055 contains thomsonite, distinguished by a small feature at 2.17 μm. There is also a 2.4 μm feature in the sample not characteristic of thomsonite but typical of many other zeolites or other hydrated mineral phases (Cloutis et al., 2002).

Finally, sample wtrfall016 has spectral features consistent with hydrated silica

(Figure 5). Opaline silica and hydrated (H2O-bearing), hydroxylated (OH-bearing) glasses exhibit absorption bands due to combination tones of Si-OH at 2.21-2.22 μm due to isolated Si-OH and at 2.26 μm due to H-bound Si-OH (Anderson and Wickershiem,

1964; Stolper, 1982; Goryniuk et al., 2004). These result in a feature near 2.2 μm that is clearly wider than the Al-OH absorption in aluminum phyllosilicates (Milliken et al.,

2008). The bands at 1.91 μm are due to H2O in the mineral structure whereas those near

1.4 μm result from both H2O and structural OH. All three absorptions can become relatively narrow as H2O is removed and only Si-OH remains (Anderson and

Wickershiem, 1964; Swayze et al., 2007; Milliken et al., 2008), but the bands in this sample are wide, indicating substantial incorporation of water in the silica structure. 217

TIR

Bulk rock samples

The infrared spectra are strongly affected by textural parameters and overlapping of multiple absorptions is common. All bulk rock samples (Figure 7; Table 1) exhibit the

H2O bending fundamental near 6 m indicating the presence of bound H2O. In the finer grained samples there is transparency feature 6-9 m, and the strong spectral rolloff has served as an indicator of fine particles in remote observations (Ruff and Christensen,

2002). The Christiansen feature (the emission maximum on the short wavelength side of the Si-O stretching minimum) occurs between 1170 cm-1 (Icel009) and 1300 cm-1

(Hvalfj054), indicating greater and lesser amounts of Si-bonding, respectively (Walter &

Salisbury, 1989). With the exception of Hvalfj054, the samples lack the box-like structure (two minima or inflections) in the reststrahlen feature near 10 m (1000 cm-1) that is characteristic of basalt and instead have a shape more typical of andesites, altered basalts, or basaltic glass (Wyatt et al., 2001; Wyatt & McSween, 2002). Hvalfj054 has this characteristic two minima reststrahlen feature and prominent features due to pyroxene. Hvalfj017 has a different reststrahlen feature, typical of Si-rich glasses or clays with a minimum at 1057 cm-1 (Wyatt & McSween, 2002; Michalski et al., 2005).

The minima near 530 cm-1 and 465 cm-1 (Figure 7) are caused by vibrations due to Al-O-Si deformation and Si-O-Si deformation with the positions on the bands shifting depending on the metal cation and mineral structure (Bishop et al., 2002b; Michalski et al., 2005 and references therein). Because these features occur paired or singly in clays and zeolites, they have been used in mapping of alteration minerals on Mars (Ruff &

Christensen, 2007). For Hvalfj017, these band positions at 531 cm-1 and 466 cm-1 most 218

closely match those of a dioctahedral ferruginous smectite (Bishop et al., 2008b) rather than endmember compositions montmorillonite or nontronite (Bishop et al., 2008b;

Michalski et al., 2005). The only other sample to clearly exhibit these two features is

Hvalfj011 at 523 cm-1 and 468 cm-1, positions which closely match terrestrial sample

Swy-1, a montmorillonite with Fe3+ substitution (Bishop et al., 2002b; Michalski et al.,

2005; Bishop et al., 2008b). Icel010 may have a weak Fe-O-Si deformation at 509 cm-1.

Icel009, Icel010 and hvalfj025 may have weak Si-O-Si deformations at 461-468 cm-1, the position found in Al- or Mg-smectites. No features in this region are apparent in

Hvalfj054.

At 630 cm-1, a feature whose cause is unattributed but appears in Al-smectites

(Bishop et al., 2002b) is seen in hvalfj054 and hvalfj011. Interestingly, no features due to metal-OH bands are apparent. These would be expected to occur from 750-950 cm-1 in clays. Particle size effects are apparent, however. The H2O bend emissivity maximum decreases in magnitude and the Si-O stretch increases in strength in the coarser size separates.

Vesicle and fracture fill

For the veins and fracture fill samples (Figure 8), Hvalfj023 exhibits most of the emissivity minima and maxima characteristic of a celadonite sample measured in Bishop et al. (2008b). Near 10 m, Hvalfj057 shares the TIR spectral features of analcime as measured by Ruff (2004). Hvalfj055 shares some characteristics with a thomsonite sample measured by Cloutis et al. (2002) although the relative strengths of features near

10 m are different. All three of these samples exhibit the H2O bending fundamental near 6 m indicating the presence of bound H2O (Figure 8). Paradoxically, the hydrated 219

silica sample does not exhibit this emissivity maximum and instead exhibits all the typical absorption features of quartz (Farmer, 1974).

XRD

Results from XRD analysis (Figure 9, 10) confirm the VNIR findings of alteration minerals in all samples.

Bulk rock samples

Hvalfj017 is the most altered of all the samples with nearly 80% smectite. Albite and hematite are present at ~10% each and the zeolite epistilbite is present in amounts

<5%. Rietveld refinement showed that the hematite is of very fine crystallite size (~80 nm) and examination of the 060 reflection near 2=62º shows that the smectite present is dioctahedral and aluminous, e.g. montmorillonite.

Other bulk rock samples were dominated by the primary mafic components plagioclase and pyroxene (Table 3). Weight percent smectite ranges from 11-22%. Minor amounts of zeolite (<10%) are sometimes present. Hvalfj025 contains ~10% smectite,

~50% plagioclase feldspar, and ~40% high-calcium pyroxene, and has the least smectite by weight of bulk rocks sampled. Hvalfj011 contains 15-20% smectite, ~55% high- calcium pyroxene, and ~25% albite. Minor amounts of the zeolite clinoptilolite (~5%) are also present. Hvalfj054 contains 15-20% smectite, 55% augite, and minor amounts of albite, hematite, levyne, and stilbite. Unusually for the bulk rock samples collected, the zeolites are a substantial component, ~10%. Icel009 and Icel010 are nearly identical with

20-25% smectite, 60% plagioclase feldspar, and 20% high-calcium pyroxene. Peak broadening suggests that icel009 may contain a small amount of amorphous silica. These samples contain non-zero but minimal hematite (<1%). 220

In hydrothermal or diagenetic systems, over time at elevated temperatures

(>40°C) in the presence of water, smectites may transform to illites or chlorites (Frey &

Robinson, 1999). They do so in a step-wise process with interstratified mixed layer clays

(illite-smectite, chlorite-smectite) as intermediate products. Modeling of the patterns from glycolated samples with the NewMod program (Yuan and Bish, 2010, in press) showed little to no evidence for interstratification. Patterns are consistent with an interstratification of smectite with 0-10% chlorite and are not consistent with the presence of illitic layers.

Vesicle and fracture fill

Non-smectite alteration minerals dominate in samples extracted from vugs and fractures within the host rocks. The mineral identifications from spectroscopy (Table 1) were all confirmed, although the actual mineral assemblage typically is more diverse

(Table 3). Hvalfj023 contains ~40% celadonite, a ferrous mica, ~35% cristobalite, and

20% smectite, with minor amounts of Na-plagioclase and mordenite. Peak broadening indicates a small amount of amorphous silica may also be present. The smectite 001 peak became asymmetric after glycolation, suggesting the presence of minor interstratifications. Hvalfj055 and Hvalfj057 were sampled from different vesicles within the host rock, separated by a few cm. Hvalfj055 is dominated by a suite of zeolites that includes thomsonite, scolecite/mesolite, and stilbite/stellerite at abundances of 25-35% for each species and smectite (8%) as a minor component. (Some of the zeolite minerals are structurally similar and cannot be differentiated without further analyses). More smectite (33%) was found in Hvalfj057, which also contains 55% analcime, 11% stilbite/stellerite, and ~1% mordenite. The wtrfall016 sample, from a white precipitate 221

within a fracture, is nearly 100% silica minerals, with ~91% quartz and 3% cristobalite along with 2% albite, and <5% clinoptiloite.

Elemental Analysis

Elemental abundance data, in weight percent oxide, are given in Table 4. Oxide totals ranged from 89.6% to 98.2% with most of the loss likely due to H2O in the samples. This loss of 2-11% is comparable to that recorded in previous Iceland geochemical analyses (Franzson et al., 2008). Based on geochemistry, the samples tend to group into four categories. Examination of K (typically mobile), Fe, Al, and Ti (typically immobile) provide particular insight (Figure 11). All bulk rock samples are similar with high FeOT and TiO2 and moderate Al2O3 and K2O. Of the vesicle fill, hvalfj016 is nearly pure SiO2, 96%, and low in all other elements. Hvaflj055C and hvalfj057C are enriched in Al2O3 with low K2O and FeOT. Hvalfj023C, in contrast, has low Al2O3, high K2O, high FeOT (relative to TiO2), and high SiO2.

Discussion

In these samples from the Icelandic plateau sequence lava flows, the bulk rock composition ranges from basalt to picrobasalt (Figure 10). These may not be precisely the original compositions of the basalt because evidence from all techniques indicates these basalts have been altered. The nature of the alteration varies by sample. For most of the bulk rock samples (Hvalfj011, Hvalfj025, Hvalfj054, Icel009, and Icel010), the primary mafic minerals pyroxene and plagioclase comprise most of the sample, and alteration minerals are <25 wt. % and mostly smectites. Bulk rock sample Hvalfj0017 is distinctly different, consisting of mostly hematite and montmorillonite suggesting a different style 222

or degree of alteration rather than subsurface fluid flow under hydrothermal conditions.

Hvalfj017 was found at the boundary between two flows (Figure 2) and may represent the uppermost surface of a lava flow in which paleosol development occurred prior to emplacement and burial by the second flow. Alternatively this may have been a concentrated pathway of fluid flow with consequently more intense leaching. Zeolites, celadonite, and silica phases are concentrated in vugs or fractures and in the bulk rock, occur in <10% abundance.

The bulk rock must have been altered by selective removal of elements into the fluids but this is difficult to discern from elemental abundance data since all bulk rock samples plot in a similar location, even though mineralogic analysis shows these samples have experienced different degrees of alteration (Figure 11). Examining the nature of the trends of Al, K, and Fe relative to comparatively immobile Ti (in primary mineral phases) in materials filling vugs and fractures vs. bulk rock samples, it is clear that different fluid chemistries or thermodynamic conditions variously favored transport and then precipitation and enrichment of K and Fe together (hvalfj023), enrichment of Al and Si

(hvalfj057, hvalfj055) or enrichment of Si only (wtrfall016). Such differences have also been reported in other Icelandic hydrothermal systems (Franzson et al., 2008).

Interestingly, the most altered bulk rock sample Hvalfj017 does not plot significantly differently from less altered samples, suggesting the alteration may have been nearly isochemical. In a ternary plot, which has been used to study weathering trends in Mars samples (Figure 12; Hurowitz & McLennan, 2007), even for the 80% smectite Hvalfj017 with its nearly complete lack of primary minerals, there is little progression along the typical terrestrial weathering trend towards the Al vertex. Most of 223

the variation lies along the feldspar-olivine join with alteration products filling veins and vesicles and becoming either Fe-enriched or Fe-depleted. Variation along this line was also typical for low water:rock ratio acidic alteration on Mars (Hurowitz & McLennan,

2007). For these particular Icelandic environments of alteration, elemental analyses of whole rock might lead to underestimation of the extent of alteration in comparison to the higher degree apparent from quantitative modal mineralogy of primary vs. secondary minerals or VNIR spectroscopy. While alteration is extensive, changes are, in a whole rock sense, largely isochemical in these samples.

All phases identified using VNIR spectroscopy were found when samples were analyzed with X-ray diffraction. Phases identified with TIR spectroscopy also yielded no incorrect spectroscopic mineral identifications. Both spectroscopic techniques positively identified smectites in Hvalfj011 (16 wt. %) and Hvalfj017 (79 wt. %), although only

VNIR techniques were able to definitively detect smectite in the other bulk rock samples

(11-22 wt. % smectite). As little as ~10% smectite is observable in VNIR spectra of altered mafic rocks measured here; however, for the XRD-determined abundances of smectite, the 2.3m absorption band depths are smaller than expected, compared to band depths from artificial particulate mixtures of basalt and nontronite (Ehlmann et al.,

2009b), possibly indicating textural effects due to differences in grain size. Zeolites in

<10% abundance also occur in the bulk rock, although their presence cannot be determined from VNIR data.

As has long been known, phases contribute to sample spectral properties out of proportion to their abundance. The presence of some alteration minerals effectively masks the presence of others. For example, analcime and celadonite in hvalfj057 and 224

hvalfj023, respectively, hide 33 wt. % and 20 wt. % smectite, respectively, in both VNIR and TIR data (Table 1, 3). In hvalfj023, the presence of celadonite (38 wt. %) completely masks the detection of silica (35 wt. %). Issues of texture and masking of some minerals by others effectively complicate spectroscopic determinations of modal mineralogy. An additional puzzle is the identification of amorphous hydrated silica phases that include Si-

OH and Si-OH bound to H2O using VNIR techniques but the identification of this same sample phase as crystalline quartz by XRD and anhydrous quartz by TIR. Although we have no explanation at present for this, it may explain the identification of hydrated silica

(Ehlmann et al., 2009) in locations on Mars where TES has identified quartz (Bandfield et al., 2004).

With the exception of sample wtrfall016, all samples contain smectite.

Interestingly, of all the techniques, VNIR spectroscopy probably provides the most information on the precise nature of the smectite-bearing component due to its sensitivity to smectite detection at small abundances as well as shifts in absorption band position that indicate the octahedral cations. Only in the case of hvalf017 could the XRD 060 peak be resolved to determine that the smectite was dioctahedral and aluminous. In other samples, interference from the diffraction peaks of other minerals prevented determination of the nature of the smectite using XRD. On Mars, the ChemMin instrument does not go to high enough 2theta angles (58°-64° required) to be able to resolve the feature. Determination of the cations Fe/Mg smectite vs. Al smectite is best accomplished at VNIR wavelengths.

225

Conclusions and Implications for Mars

In the Icelandic samples, the bulk rock is altered and smectite-bearing. Other alteration minerals such as zeolites, celadonite, and silica phases are concentrated in vugs or fractures. All phases identified using VNIR spectroscopy were confirmed using XRD analysis. As little as ~10% smectite is observable in VNIR spectra of altered mafic rocks measured here. The full extent and nature of alteration is not completely captured by spectroscopic data as textural effects and masking of some minerals’ signatures by others render some additional alteration minerals undetectable or reduce their band strength in

VNIR and TIR spectroscopy. Nevertheless, VNIR spectroscopy excellently captures the principal mineralogic diversity. Acquisition of VNIR spectra of a suite of whole rock samples from Iceland provides an accurate means of rapidly assessing the nature of past alteration. Examination of whole rocks from Mars orbit, even at large scales, similarly should provide effective information on the mineral assemblages present at smaller scales on the surface. The rocks studied here are spectrally similar to those formed at some locations on Mars, and the XRD quantitative mineralogy and geochemistry may be similar to what a rover mission to these spots would expect to encounter, with in-situ analyses. This study highlights the importance of ground-truthing spectral data from

Mars and improving techniques for quantitative remote determination of modal mineralogy but also indicates how VNIR spectroscopy provides most of the information about key mineral assemblages and, especially for smectite-bearing materials, is important supplement to information obtained in-situ by XRD and geochemical analysis.

This lends confidence to determination of past environmental conditions, including 226

hydrothermal activity and pedogenesis, on the basis of VNIR spectroscopy from Mars orbit.

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Table 1. Minerals or mineral structural elements identified by spectroscopic methods 234

Table 2. Positions of OH-related features in samples from the bulk rock XRD data analysis XRD data Table 3. 235 Combined mineralogical analysis with minerals identified using VNIR spectroscopy combined with modal mineralogy derived from Table 4. 236 Results of elemental analyses for Iceland samples(wt. %) analyses for Iceland Results of elemental 237

Figure 1. Locations of sampling sites. (a) rocks were sampled near Hvalfjordur (west) and Berufjordur (east). (b) successive basalt flows eroded by glacial activity and a stream at Berufjordur. Site of sample wtrfall016. (c) outcroping of zeolitized basalt at Hvalfjodur. At both sites, massive basaltic flows in some locations exhbited evidence of alteration minerals filling veins and fractures. 238 altered basalt from eastern Iceland. basalt. Two(e) vesicles and the bulk rock were sampled.filling types of filledwhite mineral Wtrfall016: brownish, zeolitized host rock and blue-green025: Sampledbasalt outcrops in western Iceland. (c) vesicle fill. (d) Hvajfl054, 055, 057: Hvalfj023, (red), overlain by a younger flow (gray). (b) Icel009, 010: Sample development flow with paeleosol a between represent contact a Figure 2. Study samples from weathering and hydrothermalenvironments. (a) Hvalfj011, 017: Sampledfriable units cut by a stream--may Typical fractures in s from 239

Figure 3. ASD spectra of whole rocks displayed in Figure 2. Spectra from multiple rock faces were averaged to a single spectrum. Spectra were aquired with an 18-degree foreoptic at a 10 cm standoff distance at an incidence angle of ~10° and emission angle of ~20°. 240

A B

Figure 4. Bulk rock bidirectional visible near-infrared (VNIR) reflectance for the <150 m and <25 m particle size separates from samples inferred to contain (a) Fe/Mg phyllosilicate and (b) Al phyllosilicate. 241

A B

Figure 5. Vesicle and fracture fill bidirectional visible near-infrared (VNIR) reflectance for the <150 m and <25 m particle size separates. 242

A B

Figure 6. Metal-OH related absorptions in key spectral regions for bulk rock <150 m and <25 m particle size separates as in Figure 3 for the spectral regions with (a) combination tones and (b) stretching fundamentals. 243

Figure 7. TIR reflectance spectra of bulk rock coarse (75 or 105-500 m) and <150 m size fractions, converted to approximate emission spectra using Kirchoff’s law (1-R=E; accuracy may be influenced by particle size effects—see text). Positions of key absorptions are indicated. Laboratory spectra inverted reflectance from the USGS library and emission spectra from the TES spectral library. 244

Figure 8. TIR reflectance spectra of vesicle and fracture fill coarse (75 or 105-500 m) and <150 m size fractions, converted to approximate emission spectra using Kirchoff’s law (1-R=E; accuracy may be influenced by particle size effects—see text). Positions of key absorptions are indicated. 245

Figure 9. XRD patterns from bulk rock samples 246

Figure 10. XRD patterns from vesicle and fracture fill samples 247

Figure 10. Total alkalis silica diagram for the analyzed samples computed from volatile- free weight percent oxide data 248

Figure 11. Elemental abundance data, in weight percent, for (a) Al2O3 vs. TiO2, (b) K2O vs. TiO2, (c) FeOT vs. TiO2 and (c) FeOT vs Al2O3 showing trends in mobility of cations with alteration. 249

Figure 12. Ternary Al2O3, (CaO+Na2O+K2O), (FeOT+MgO) diagrams, data plotted in mole percent. The bold black arrow indicates the direction of terrestrial weathering (after Hurowitz & McLennan, 2007). Light gray arrows indicate relationships between bulk rocks (blue diamonds for Fe/Mg smectite-bearing, green diamonds for Al smectite-bearing) and vesicle/fracture fill (orange squares) found within those bulk rocks. 250

Figure 13. Band strengths of the Fe/Mg-OH combination tone in Iceland samples are weaker than expected compared to band strength from synthetic mixtures of nontronite and basalt (45-75 m) made in the laboratory (Ehlmann et al., 2009b). 251

Chapter 7:

MODAL MINERALOGY OF HYDRATED SILICATE-BEARING

MIXTURES DERIVED FROM VISIBLE/NEAR-INFRARED

REFLECTANCE: AN ACCURACY ASSESSMENT OF HAPKE

AND SHKURATOV RADIATIVE TRANSFER MODELS

Bethany L. Ehlmann1, John F. Mustard1, Francois Poulet2

1Department of Geological Sciences, Brown University ([email protected]) 2Institut d’Astrophysique Spatiale, Université Paris-Sud

251 252

ABSTRACT—Here we report preliminary results from a test of the efficacy of Hapke and

Shkuratov radiative transfer models in determining modal mineralogy of mixtures bearing hydrated minerals from visible/near-infrared (VNIR) spectra. Bidirectional VNIR spectra were acquired of particle size separates of the endmembers, nontronite, olivine, basaltic glass, and basalt, as well as their mixtures. Initial results with binary mixtures show that sample modal mineralogy can mostly be estimated to within ~10% for mixtures of endmembers of the same grain size. Future tests of the two radiative transfer models will be conducted using mixtures and

VNIR spectra from multiple grain size mixtures with a greater number of endmembers. Results from the preliminary tests discussed here suggest that variability in the spectral reflectance of replicate samples, uncertainties in optical constant derivations, and attention to the different spectral behaviours of mixtures in different wavelength regions must be considered as testing and application of radiative transfer models to VNIR spectral data from mixtures containing hydrated materials continues.

Introduction

The discovery of diverse phyllosilicates on Mars using visible/near-infrared (VNIR) spectral data from the orbiting imaging spectrometers points to aqueous environments with long- lived water-rock interaction [Poulet et al., 2005; Bibring et al., 2006; Mustard et al., 2008;

Murchie et al., 2009]. Determining the extent and nature of alteration requires initial mineral identification, but single minerals rarely alone provide unique discrimination of formation environments. Rather, the mineral assemblages comprising rocks or sediments and the relative quantities of the various minerals, i.e. modal mineralogy, are attributes more uniquely characteristic of formation environment. Knowledge of modal mineralogy allows narrowing the 253 permissible thermodynamic conditions (T, pH, Eh, ion activities) and reaction rates and/or durations of alteration.

Volumetric abundances of coarse grain size mineral constituents in surface units are possible to determine remotely via linear deconvolution of thermal emission data to accuracies of

10-15% [e.g. Ramsey and Christensen, 1998] because the photon interactions are mostly singly scattered. In VNIR wavelengths and for finer grained particulate mixtures, multiple scattering dominates instead and spectral mixing is nonlinear, i.e. absorption band strength is not directly proportional to abundance. Up to now, VNIR spectroscopy has essentially been used for identification, not quantification of mineral phases. However, the identification of numerous diverse hydrated silicates on the surface of Mars using VNIR spectroscopy and their non- detection at TIR wavelengths necessitates new attention to the development of VNIR techniques for quantification of modal mineralogy. The objective of this research is to move from qualitative VNIR data analysis towards quantitative VNIR analysis.

Radiative transfer models predict that VNIR mixing systematics can be modeled, albeit non-linearly [Hapke, 1993; Shkuratov et al., 1999]. There have been previous attempts and successes in modelling mineral abundances for specific types of simple mixtures, e.g. clay + carbon black [Clark, 1983] and anhydrous mafic mineral mixtures. Laboratory studies have shown the Hapke [Mustard and Pieters, 1989] and Shkuratov [Poulet and Erard, 2004] radiative transfer models predict modal abundance to ~10% for well-controlled mixtures of mafic minerals. But while model estimates have been accurate for a handful of anhydrous mixtures in laboratory, their performance for determining hydrated mineral abundance in multi-component mixtures has not yet been evaluated. Because of the considerable compositional variability of hydrated minerals and changes in spectral properties related to hydration state, these alteration 254 minerals may prove more challenging to model than mafic minerals. Radiative transfer theories have been applied to the modelling of asteroid ice and organic components [e.g. Cruikshank et al., 1998] and silicate mineralogy [e.g. Clark, 1995]. For Mars data, the Shkuratov model has been applied to estimate mineral abundances within phyllosilicate-bearing terrains using

OMEGA data [Poulet et al., 2008], providing a first comparison of variations in hydrated mineral abundance from place to place. However, to gain confidence in the accuracy of these estimates from remote data, fully characterizing model accuracy over a suite of samples in laboratory is essential.

The key barriers to date for widespread implementation of these models are both theoretical (e.g. problems treating spectrally neutral species and dark (albedo<0.2) species, porosity effects [e.g. Hapke, 2008]) and practical (lack of well-characterized mixtures of known mineralogy and grain size to test the models). We will address the latter, by generating spectral data of well-defined mixtures, using best-fit techniques and the Hapke and Shkuratov radiative transfer models to estimate mineral abundances in these mixtures, and assess the accuracy and precision of model results with hydrated minerals. We address the former by evaluating any patterns in model inaccuracies. It may also be possible to determine, e.g. via sensitivity analyses, systematic trends due to error in the models’ underlying assumptions and propose improvements.

We report here the results from a pilot test of the efficacy of both the Shkuratov and

Hapke models with hydrated minerals. After briefly outlining the theory and principal differences between the two radiative transfer models, we present the spectra acquired of endmember particle size separates and prepared mixtures. We then discuss mixing results from same-grain-size binary mixtures of nontronite (an Fe smectite clay) with olivine and basaltic glass as well as mixtures of nontronite with basalt (producing a 5 component assemblage). We 255 show spectra and discuss future work involving ternary mixtures, multiple grain size combinations, multiple hydrated minerals, and potential refinements to raditative transfer models.

Methods and Materials

Endmembers and mixtures

Our first experiments mix a hydrated phyllosilicate with a relatively bright, transparent component and relatively dark, opaque component (Table 1). Nontronite, a dioctahedral smectite clay with iron as the major octahedral cation, was obtained from the Clay Mineral Society, source clay NG-1 [Schneiderhohn, 1965; Lear et al., 1988]. The bright, relatively transparent material chosen was olivine (~Fo90) from San Carlos, Arizona, which was handpicked to remove pyroxenes and grains with dark inclusions. The dark, homogenous component was fresh, anhydrous basaltic glass, obtained from the Big Island of Hawaii. An additional endmember was dark and a mixture of multiple minerals, basaltic rock sample 79-3b from Medicine Lake, which had already been well-characterized by Wyatt et al. [2001]. Point count data of 79-3b, accurate to within ~5%, show that it contains 76% plagioclase, 8% high-calcium pyroxene (augite or diopside), 5% pigeonite, 5% Si-K2O glass, 1% orthopyroxene, and 1% olivine [Wyatt et al.,

2001]. Samples were measured with XRD to verify composition (Figure 1). There were small amounts of plagioclase in the basaltic glass and quartz in the nontronite but samples were otherwise relatively pure.

The samples were crushed and sieved into eight particle size fractions (<25 m, 25-45

m, 45-63 m, 63-75 m, 75-106 m, 106-125 m, 125-150 m, 150-250 m, 250-500 m). The amount of quartz in the NG-1 sample was found to be relatively coarse-grained and mostly 256 removed by sieving, so the largest NG-1 endmember size fraction is 75-106 m. Nontronite was dry sieved while olivine and basaltic endmembers were wet sieved. From these endmembers, a suite of mixtures were produced, with fractions given by weight. Binary mixtures at 45-75 m were made of the endmembers at 5%, 10%, 30%, 50%, 70%, 90% for the initial testing. More complex ternary and multi-grain size assemblages were also prepared and measured (Table 2).

Measurements

Reflectance spectra were measured in the NASA/Keck Reflectance Experiment

Laboratory (RELAB) bidirectional spectrometer (Pieters, 1983). Samples were placed in dark sample holders with a diameter of 1 cm or greater and a depth of at least 3 mm. Larger and deeper dishes were used for more transparent materials, e.g. coarse olivine particle size separates, to ensure that the bottom of the dish was not contributing to the light on the sensor.

Unpolarized monochromatic light was directed onto the uncovered horizontal sample at an illumination angle of 30º. The reflectance of the sample and a pressed halon reference standard were measured in turn from a rotating sample platform for each 5 nm spectral increment at an emission angle of 0º. Data were acquired from 0.28-2.6 μm. These first analyses focus on the

1.0-2.6 μm region that is covered by the infrared detectors on CRISM and OMEGA.

Radiative transfer modeling approach

The two most important factors determining the VNIR spectrum of a sample or surface are composition and grain size. Two principal radiative transfer models exist for VNIR light interacting with particulate mixtures, the Hapke model [Hapke, 1981; 1993] and the more recently promulgated Shkuratov model [1999]. There are differences in the fundamental physical assumptions underlying the two models. In Hapke modeling, at present the more commonly 257 employed of the two, surface reflectance, r, reduces to a function of incidence (i) and emission

(e) angle and single scattering albedo, w, over e ~ 0 º and i = 10-40º. That is,

(1)

where Chandrasekhar’s function for isotropic scattering is H(μ)=(1+2μ)/(1+2μ(1-w)1/2). This is a simplified version of Hapke’s model in which phase functions are neglected and particulate surfaces are assumed to be isotropic scatterers of light (e.g. Mustard & Pieters, 1987). Single scattering albedo, w, is the ratio of scattering efficiency to extinction (scattering + absorption) efficiency and varies in a characteristic manner as a function of wavelength for surfaces of different composition and grain size. For a multi-component particulate mixture, the Hapke model assumes wmix varies linearly as sum of wi for each of the components weighted by its cross-sectional area, which can be expressed in terms of measurable parameters such as mass

(Mi), density (i) and particle diameter (Di). Given a database of VNIR spectra of minerals, ri for each mineral can be converted to wi , using a lookup table. Least squares inversion or more complex deconvolution techniques can then be utilized to solve for the fraction of each, Fi, which can then be converted to mass fraction Mi when i and Di are known since the relationship is expressed as

M i n N D i i i i (2) wmix B Fi wi where Fi B N M i 1 i i B i i Di

For the results covered here, fractional values sum to one for each sample.

The Shkuratov model for estimating modal mineralogy is more complex, involving probabilistic treatment of forward and backward scattering through particles, modeled as a 1-d stacking of slabs [Shkuratov et al., 1999]. It relies on having, not endmember reflectance spectra, 258 but wavelength dependent endmember optical constants n and k, the refractive index and extinction coefficient. (Note: Hapke modeling too can employ n and k but the mixing systematics are linearized with just w). The optical constants are either obtained through transmission spectroscopy or derived using spectra from multiple particle size separates of minerals [Roush et al., 2007]. In this case, optical constants of endmembers were estimated by using the Shkuratov model to iteratively determine the imaginary index of refraction for each endmember from reflectance spectra of the endmember particle size separates [Poulet & Erard, 2004; Roush, et al.2007]. The advantage of the Shkuratov model is that it allows simultaneous calculation of endmember abundance and particle size. The Shkuratov model employed here, adapted from

Poulet and Erard [2004], uses downhill simplex methods to select endmember parameters, minimizing error between measured and modeled spectra.

Initial Results

Endmember particle size separates

Figures 2-5 show the endmember particle size separates and Figure 6 shows the 45-75

m size separate for each of the endmembers, permitting comparison of their relative reflectances at different wavelengths. Nontronite has the most complex suite of absorptions consisting of Fe-OH and H2O vibrational overtones and combination tones at 1.43, 1.91, 2.29,

2.40, 2.50 m as well as electronic absorptions due to Fe3+ at 0.45, 0.66, and 0.97 m (Frost et al., 2002; Bishop et al., 2002). Olivine has a small feature at 0.64 m and deep broad overlapping electronic absorptions centered near 1.05 m. Basaltic glass has broad electronic absorptions centered at 1.07 m and 1.96 m, and basalt endmember has an electronic absorption centred at

0.98 m and what is perhaps a shallow, broad feature in the 1.9-2.6 m range. In general, all 259 spectra have a rapid dropoff in reflectance shortward of 0.6 m. The anhydrous endmembers have relatively flat or slightly increasing reflectance longward of 1.5 m, while the hydrated nontronite endmember declines markedly in reflectance in this spectral region. The spectral contrast of olivine and nontronite is high, while the dark (R < 0.2) basaltic endmembers have noticeable features but small reflectance contrast in absolute terms.

In general, all endmembers show decreasing reflectance as particle size increases. An exception is the <25 m particle size separate for olivine (Figure 3), which is considerably darker than coarser-grained size separates at wavelengths great than 1.5 m. This result is also consistent with findings by Pieters [1983], indicating that for wavelengths of very high reflectance the trend can reverse, i.e. bright materials become darker with decreasing particle size. Absorption band depths increase in relative strength as particles become coarser, which is apparent in the normalized reflectance data for these samples (not shown).

Binary, single particle size mixtures

Spectra of nontronite-olivine and nontronite-basaltic glass for the 45-75 m suite are shown in Figures 7-8. As expected, phyllosilicate vibration band strengths decrease with decreasing phyllosilicate abundance. This trend is more pronounced in the glass mixtures where dark components cause a lower apparent band strength compared to the olivine mixtures for equivalent phyllosilicate abundance, e.g. in the diagnostic 2.29 m Fe-OH combination tone.

Low abundance mixtures of nontronite with the basaltic glass show some irregularities in the continuum reflectance (e.g. 10% nontronite mixture is darker than expected; Figure 8) perhaps related to inhomogeneous distribution of nontronite in the sample dish. In the olivine samples, phyllosilicate absorptions are quite pronounced near 2.29 m with even as little as 2 wt. %, while 260 the 1.4 m absorption within the broad olivine absorption is not apparent until 30 wt. % nontronite.

Mixtures were also prepared of the olivine and basaltic glass endmembers (Figure 9). As in the nontronite-olivine mixtures, the influence of the second component in the 1.5-2.6 m region is quite apparent. The broad 2 m absorption of basaltic glass is visible even at only 5-

10% abundance. As is also the case for the nontronite-olivine mixtures, the influence of olivine is difficult to discern when its abundance is less than 50%. The last suite of binary, single particle size mixtures prepared was that of nontronite with basalt 79-3b (Figure 10). The magnitudes of the phyllosilicate vibration bands are comparable to those in the nontronite- basaltic glass mixtures.

Ternary mixtures were prepared of nontronite, basaltic glass and olivine (Figure 11).

These display much more widely varying continua shapes and absorption band depth variations.

Additionally, mixtures were prepared of 45-75 m nontronite with coarser-grained (125-250 m) olivine (Figure 12) and basalt (Figure 13). For the same abundance of nontronite, mixtures with coarser grained olivine are darker, but also exhibit stronger phyllosilicate absorption band depths

(Figure 14). With the basalt, the mixtures including coarser grains are either of comparable brightness or quite a bit darker than the 45-75 m mixtures, but for the same abundance of nontronite, the phyllosilicate absorption band depths are generally weaker (Figure 15). Indeed for the 5% nontronite mixture, no nontronite-related bands are apparent when mixed with coarse basalt. These coarse anhydrous-medium-sized hydrated mixtures were also measured in triplicate to determine the inherent variability in reflectance due to random grain orientations and effects of sample pouring into the dish (Figure 12, 13). There is sometimes overlap between size fractions in terms of overall albedo, especially for the darkest mixtures (e.g. Figure 13b). 261

Hapke mixture modelling – first results

Radiative transfer modelling has thus far only been performed on binary mixtures of nontronite with other minerals at the same grain size (45-75 m). For the nontronite-olivine mixtures, the Hapke model performs well. The modeled fnont is ±10% of actual for all of the mixtures when actual fnont is greater than 10%, and RMS error is <0.005 (Figure 16a). For fnont

<10%, fitting the small magnitude vibrational absorptions due to nontronite in the very bright part of the spectrum (1.6-2.6 m) is sacrificed by the model’s fitting the high contrast 1 m olivine band (Figure 17a), resulting in underestimation of the amount of nontronite.

The Hapke model has greater difficulty correctly modelling mixtures containing darker materials. For the basaltic glass-nontronite mixtures, RMS error is <0.015 and fnont is within 12% of actual (Figure 16b). Selectively modeling only using certain wavelength ranges shows that poor fits in the highest contrast portion of the spectra from 350-600 nm results in decreased fit accuracy, especially for samples with approximately equal amounts of volcanic glass and nontronite. Nontronite is overestimated. Fits and modeled spectra for small fnont samples are extremely good, however (Figure 17b). A similar pattern to the basaltic glass mixture, albeit less pronounced, is observed in the mixtures with basalt (Figure 16c, 17c).

Sensitivity analyses were conducted to examine how results varied with wavelengths under consideration, changes to input parameters, and model endmembers. Tests were done comparing the results from Hapke modelling using Chandrasekhar’s function as formulated in

Hapke (1993) to more recent recalculated H functions from Hiroi (1994). Differences in modeled mineral abundances and RMS varied by <1%, so H function formulation is judged to be of negligible influence for these mixtures. Tests were also run including all four endmembers for which optical constants were available instead of just the two actually present in the mixture. The 262 model was robust to inclusion of extra endmembers and did not include extraneous minerals in fits at >10% levels. However, the dark materials remained systematically undermodeled by

~10%.

Shkuratov mixture modelling – first results

Radiative transfer modelling has thus far only been performed on binary mixtures of minerals at the same grain size (45-75 m). RMS errors in fit between measured spectra and modeled spectra are less than 0.01 for nontronite-basaltic glass, nontronite-basalt, and basaltic glass-olivine mixtures. Modelling accurately predicts fractional weight percents of constituents to within ~10% (Figure 18). Grain size estimates vary more widely, falling within the 45-75 m actual size range of the particles for the mixtures of basaltic glass with nontronite and olivine, and are somewhat overestimated for the nontronite-basalt mixture.

Interestingly, estimates of modal mineralogy and grain size in the olivine-nontronite mixture are poor. This is attributable, however, to the optical constants used. With initial tests using library optical constants instead of those derived from the endmembers themselves, fnont was modeled to within 10% of its actual abundance [Ehlmann et al., 2009]. Subtle differences in the derived vs. library optical constants are shown in Figure 19. In particular, differences in the narrowness and shape of the phyllosilicate absorptions and the magnitude of the olivine feature may be responsible for differences in spectral fits.

An additional difficulty with optical constants is that the Shkuratov model does not permit accurate modelling when R < 0.2 (Poulet, personal comm.). For derivation of the basaltic glass optical constants, only samples with grain size >63 μm were used. Shkuratov modeling is not currently possible for the coarse grained basalt with nontronite mixture because the samples with 50% or less nontronite are too dark. 263

A final preliminary test of the Shkuratov model was in estimation of mineral abundances within BAS 79-3b, the polymineralic basaltic endmember with well-constrained mineralogy.

Library optical constants of common minerals found in basalt were used in the fits. Assuming point count data in Wyatt et al. [2001] are accurate, the Shkuratov model accurately estimates to within 10% modal mineralogy for the BAS 79-3b grain size fractions that are >100 μm. When modeling was done using the finer grained fractions of the endmembers, however, the mineral assemblage modeled changed and accuracy decreased (Figure 20).

Discussion

Qualitative examination of the VNIR spectra of endmembers and their mixtures provides a number of interesting insights relevant to the spectral analysis of phyllosilicate-bearing mafic materials on Mars. Dark, coarser grained mafic materials appear to be more effective at “hiding” the sharp vibrational absorptions of hydrated minerals. As shown in Figure 15, at the 5% level it is difficult or impossible to visually discern the presence of phyllosilicates. Future work will test whether this is also the case for quantitative radiative transfer modeling. This has important implications for Mars where much of the surface is dark basaltic rock, and bedrock stratigraphies are most useful in reconstructing environmental histories. The coarser grained mixtures with mostly transparent olivine or fine sand 45-75 m mixtures with equivalently sized phyllosilicate are probably less relevant to deciphering Mars stratgraphic sections, but in these cases phyllosilicate absorptions are more prominent. Interestingly, for olivine, phyllosilicate absorptions are deeper in mixtures with coarse olivine rather than fine, presumably due to the fact that most light passes through the coarse olivine and is not scattered until phyllosilicates are encountered. 264

The primary purpose of these mixtures is, however, to assess the prospects for quantitative determination of abundance from VNIR spectra. Using both the Shkuratov and

Hapke models, endmember abundance is mostly predicted within ±10% with low RMS error for binary mixtures. This is not universally the case, however, and each model exhibits difficulty with different mixtures. Hapke model errors are sometimes greater than 10%, especially for the mixtures of nontronite with darker materials, basalt and basaltic glass. These are underestimated, relative to nontronite. The underestimation of dark components here is similar to the underestimation of dark materials (magnetite with olivine) also reported by Mustard & Pieters

(1987) with Hapke modeling. High contrast wavelength regions of the spectrum also generate poor fits, e.g. the olivine 1 μm absorption and the decrease toward visible wavelengths in dark materials. Shkuratov model results are fairly accurate with the dark endmembers, but poor with the bright olivine.

These preliminary findings have revealed a number of avenues that must be pursued to understand the accuracy of laboratory data and limitations of the two radiative transfer models so as to gain confidence that remotely sensed data can be understood and modeled. For the

Shkuratov model, a key question is variation in optical constants and what degree of accuracy in their derivation is necessary for robust modeling. The choice of mineral optical constants influences the modeled mineralogy for the nontronite-olivine mixtures in a way that is not completely understood. At present, optical constants derived from the endmembers according to the methods of Poulet and Erard (2004) and Roush et al. (2007) do not perform as well as library optical constants. A second puzzle concerns the effects of grain size, both in the proper allowable range of model grain size variation and how this influences modal mineralogy determination as well as how measurements of different grain-sized materials apparently lead to different modal 265 mineralogies (Figure 20). Additionally, the Shkuratov model is formulated to deal with materials that are relatively transparent. Shkuratov et al. (1999) do not evaluate k values above a few tens of 10-3 and Shkuratov and Grynko (2005) explicitly state that the imaginary part of the refractive index of particles is assumed to be small. Samples with R~0.1, as in the mixtures with coarse basalt, may prove problematic; yet, to date, the Shkuratov model appears to work better on dark materials rather than bright.

For the Hapke model, a full implementation using n and k should be trialed. Additionally the performance with dark materials and in high spectral contrast wavelength regions is less than ideal. Yet the Hapke model has so far been shown to be robust against the use of different H functions as well as the inclusion of extraneous endmembers. Incorporating porosity parameters

(Hapke, 2008) and perhaps different phase functions for different constituents (Mustard &

Pieters, 1989) may prove useful to refine this model’s implementation, should further inaccuracies be found following the modeling of VNIR spectra from additional mixtures.

In future studies, both radiative transfer models will be assessed with all of the mixtures prepared as well as with more complex mixtures involving multiple hydrated components and

N>3. As recent results from coarse grained mixtures show, experimental errors or variability in spectral measurements may also contribute to radiative transfer model error. Especially for darker mixtures, the amount of spectral variability is comparable to the variation due to mineral abundance when samples are measured in triplicate (Figure 12, 13). This is probably due to the heterogeneous nature of the samples, random grain orientation effects upon pouring into the sample dish and, perhaps, some segregation of particulates due to density differences. Going forward, all mixtures will be measured in triplicate so that the effect of spectral variability of 266 mixtures can be translated into variation in modeled modal mineralogies and errors due to sample variability can be separated from errors due to inherent model limitations.

Summary and Future Work

Pure endmember particle size separates have successfully been prepared of four endmembers, including hydrated nontronite. These have been used to produce mixtures of well- constrained composition, relevant to the interpretation of VNIR spectral data from Mars. Initial pilot tests of the application of Hapke and Shkuratov radiative models on VNIR spectra from binary mixtures of nontronite with equivalent sized olivine, basaltic glass, and basalt show that modal mineralogy and the abundance of phyllosilicates can be calculated to within 10% of actual abundances with acceptable RMS error most of the time.

However, several potential challenges were also noted. The grain size predictions allowed by the Shkuratov model do not successfully predict the actual grain size of the constituents, and the modal mineralogy of the basalt endmember was not successfully predicted from existing optical constants. Optical constants derived from VNIR spectra of these endmembers disagree somewhat from library optical constants when these exist for comparison, and these can profoundly affect determination of modal mineralogy. The Hapke model showed substantially higher error in the visible wavelengths. Additionally, measured spectra of the mixtures were sometimes found to have a fairly high degree of variability, which may affect model results.

Future work will address these challenges and others. All mixture spectra will be remeasured at least twice to determine the sensitivity of modal mineralogy determinations to variability in spectral properties resulting from sample preparation or random orientation effects. 267

Both radiative transfer models will be tested with more complex mixtures (N>3) with multiple grain sizes, for which some spectra have already been acquired. Future mixtures will include neutral species, multiple hydrated materials, and perhaps samples derived from whole rocks so as to test model efficacy under realistic geologic conditions. We expect future work will fulfill the research objective of moving VNIR spectral analysis from the qualitative endeavor of identification towards quantification of mineral abundance

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instrument on MRO, Nature, 454, 305–309.

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oxides in nontronite from Hohen Hagen, Federal Republic of Germany. Clays and Clay

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Poulet, F, Erard, S. (2004) Nonlinear spectral mixing: Quantitative analysis of laboratory

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Table 1. Endmembers used in mixtures. Material Est. density Endmember Source/Description Characteristic (g/cm3)

Nontronite hydrated 2.2-2.3 Clay Minerals Society source clay from Hohen Hagen, (NG-1) mineral Germany. Nontronite within coarse quartz sandstone with some iron oxides

Olivine bright, 3.27-3.37 ~Fo90 olivine from San Carlos Reservation, Arizona (OLV-SC) transparent Basaltic glass dark, 2.7-2.85 Unaltered basaltic glass sample (aphanitic) from the Big Island, (BAS-Gl) homogenous Hawaii (from M. Wyatt)

Basalt dark, multi- 2.7-3.2 Feldspar, pyroxene dominated basalt from Medicine Lake, (BAS 79-3b) mineral Oregon provided by M. Wyatt; modal mineralogy described in mixture Wyatt et al., JGR, 2001

Table 2. Mixtures prepared. # Endmembers Particle size Constituents Binary 45-75 m NG-1, OLV-SC Binary 45-75 m NG-1, BAS-GL Binary 45-75 m OLV-SC, BAS-GL Binary* (5 minerals) 45-75 m NG-1, BAS 79-3B (76% plag., 8% HCP, 6% LCP, 5% glass) Ternary 45-75 m NG-1, OLV-SC, BAS-GL Binary 45-75 m, 125-250 m NG-1 (med.), OLV-SC (large) Binary 45-75 m, 125-250 m NG-1 (med.), BAS-GL (large)

272 Figures

Figure 1. XRD patterns from the four endmembers. The nontronite contains small amounts of quartz and the basaltic glass small amounts of plagioclase, but samples otherwise have their expected compositions. 273

Figure 2. VNIR spectra of particle size separates of the nontronite (NG-1) endmember 274

Figure 3. VNIR spectra of particle size separates of the olivine (OLV-SC) endmember 275

Figure 4. VNIR spectra of particle size separates of the basaltic glass (BAS-GL) endmember 276

Figure 5. VNIR spectra of particle size separates of the basalt (BAS 79-3b) endmember 277

Figure 6. VNIR spectra of endmembers for the 45-75 m particle size separate 278

Figure 7. VNIR spectra of mixtures of OLV-SC and NG-1, each at 45-75 m 279

Figure 8. VNIR spectra of mixtures of BAS-GL and NG-1, each at 45-75 m 280

Figure 9. VNIR spectra of mixtures of BAS-GL and OLV-SC, each at 45-75 m 281

Figure 10. VNIR spectra of mixtures of BAS 79-3b and NG-1, each at 45-75 m 282

Figure 11. VNIR spectra of ternary mixtures of BAS-GL, OLV-SC and NG-1, each at 45-75 m 283

Figure 12. VNIR spectra of binary mixtures of OLV-SC and NG-1, at 125-250 m and 45-75 m, respectively. Replicate measurements are shown as dashed lines. 284 measurements are shown as dashed lines. (b) zoom in on low reflectance measurements in on low reflectance lines. (b) zoomshown as dashed are measurements Figure 13. (a) A VNIR spectra of binary mixturesand 45-75 m,respectively.of BAS 79-3b and NG-1, at 125-250 m Replicate B 285 A (b) spectra of the 1700-2500 region showing nmcontinuum-removedband depths Figure 14. (a) VNIR spectra of binary mixtures of 45-75 m NG-1 with 125-250 m (solid lines) or 45-75 m (dashed lines). B 286 A continuum-removed spectra of the continuum-removedregion showing band depths 1700-2500 nm Figure 15. (a) VNIR spectra of binary mixtures of 45-75 m NG-1 with B 125-250 mm (solid lines) or 45-75 (dashed lines). (b) 287 wavelength regions and the lower plot gives RMS mixtureerror for each in the spectral fit. Fits percent abundance were done utilizing differe Figure 16. eut rmHperdaietase oeig h pe ltcmae oee otoiet culnnrnt weight nontronite to actual transfer modeling.nontronite radiative modeled The upper plot compares Results fromHapke nt 288 nmmes(magenta). endmembers Figure 17. NRsetamdlduigHperdaietase oeig h niigruieue ihr2edebr (blue) or 4 2 endmembers routine used either transfer modeling. The unmixing radiative using Hapke VNIR spectra modeled 289 A B

C D

Figure 18. Results from Shkuratove radiative transfer modeling. The upper plot compares modeled nontronite to actual nontronite weight percent abundance, the middle plot gives the RMS error in the spectral fit, and the lower plot gives the actual (dashed line) and modeled particle sizes for (a) NG-1 and OLV-SC, (b) NG-1 and BAS-GL, (c) NG-1 and BAS 79-3b, and (d) BAS-GL and OLV-SC. 290

A B

Figure 19. The absorption coefficient, k, derived from the endmembers (black) and from literature optical constants (blue) shows discrepancies for both (a) olivine and (b) nontronite. 291

Figure 20. Shkuratov model-derived optical constants for the different particle size separates fo BAS 79-3b compared to the point count weight fraction determined in Wyatt et al. (2001) (large black diamonds). 292

Chapter 8:

CONCLUSIONS

292 293

The differences between Earth today and Mars today are striking. Mars has a thinner atmosphere of mostly carbon dioxide and no active hydrologic cycle. Mars does not have, and appears never to have had, plate tectonics. Mars’ nearly uniform basaltic composition contrasts with the diverse granitic, basaltic, and sedimentary units found on

Earth.

Yet, recent discoveries of alteration minerals in ancient crust from both orbital and landed missions point to the fact that ancient Mars was also incredibly different from

Mars today. The past five years have been a period of rapid discovery in terms of mineralogical indicators of past Martian environments with liquid water. The 20 different types of alteration minerals found to date on the surface of Mars are listed in Table 1.

Some occur in assemblages with distinctive geomorphic settings or in association with other minerals (e.g. Chapter 3); but some occur in multiple settings that are more difficult to decipher (e.g. Chapter 4). Nevertheless, a key finding from synthesizing this new mineralogic record is that one must think locally, not globally, to understand the aqueous chemical environments of Mars.

The key advance of high resolution VNIR imaging spectroscopy has been, not only the discovery of new mineral classes, but the contextualization of mineralogy with morphology at high resolution. Like present-day Earth, ancient Mars apparently hosted diverse geologic environments, which varied in space and in time. Below, I synthesize findings concerning the nature of aqueous environments during Mars’ first billion years, discuss the key outstanding questions arising from this work, and offer some thoughts as to how these findings influence strategies for future Mars exploration.

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The nature of aqueous processes on earliest Mars: the first billion years

With the abundance of new data on the mineralogic diversity of Mars, the general paradigm of Bibring et al. (2006) has held: the earliest rock record preserves evidence for aqueous alteration under neutral to alkaline conditions while subsequent (but still ancient) periods typically preserve evidence for aqueous alteration under acidic, often water- limited conditions. There appears to be a clear, unidirectional trend in Mars’ aqueous chemical evolution that occurred during the first billion years. By ~3 Gyr, or the late

Hesperian/early Amazonian, there is still evidence for intermittent fluid flow in the case of a few, later-active valley networks (Fasset & Head, 2008), but post-Hesperian rocks do not preserve chemical evidence for alteration comparable to that from rocks of the first billion years of Mars history.

A synthesizing effort by the CRISM science team has led to the characterization of Mars’ distinctive aqueous chemical environments based on their alteration mineralogy and geologic setting. The time-ordering of these environments, from the work of Murchie et al. (2009) and with modifications from this and subsequent work, is shown in Figure 1.

Phyllosilicates in fans, plains sediments (i.e. chlorides; Osterloo et al., 2008;

Wray et al. 2009), intracrater clay-sulfates, and Meridiani layered deposits all represent the products of deposition in standing bodies of water of varying pH. The systems ranged from open to closed basins and from feeding by an extensive watershed, as in the case of

Jezero crater (Fasset & Head, 2005; Ehlmann et al., 2008a), to ephermeral upwelling of groundwaters in sabkha environments as in the sedimentary units at Meridiani Planum

(Grotzinger et al., 2005; McLennan et al., 2005; Arvidson et al., 2006). While all of these systems involve surface waters, the nature of the aqueous geochemistry varied

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substantially. Carbonate/smectite assemblages that indicate neutral to alkaline waters are found in Jezero crater (Ehlmann et al., 2008; 2009), sulfate-Fe/Mg smectite assemblages are found in Gale crater (Milliken et al., 2010), and Al-sulfate/clay assemblages that form in acid saline lakes are typical within the craters of Terra Sirenum (Wray et al., 2009;

Swayze et al., 2008; Baldridge et al., 2009). Groundwater certainly played a role in many of these systems, but deposition of the mineral bearing units was at the surface, implying stable or metastable waters on the surface of Mars from at least the late-Noachian to

Hesperian.

Additional evidence for the influence of surface or near-surface waters is provided in intact stratigraphies of rock units bearing different alteration minerals: layered phyllosilciates and carbonate deposits. In and around Mawrth Vallis (McKeown et al., 2009; Loizeau et al., 2010), in Terra Meridiani (Wiseman et al., 2008), and in and around the Nili Fossae (Ehlmann et al., 2009) aluminum phyllosilicates are found overlying Fe/Mg smectites in stratigraphies extending over hundreds to thousands of kilometers. Additionally, in the Nili Fossae region, carbonate rocks associated with an olivine-rich unit are found overlying Fe/Mg smectites in an equally regionally extensive stratigraphy. The most plausible formation environment to explain these recurring units is one of surface to near-surface weathering, where the final products are controlled by the chemistry of the precursor rocks. For typical mafic materials, additional loss of Fe, Mg and Ca cations occurs preferentially during leaching. Remaining products, relatively enriched in Al and Si, include kaolinite, montmorillonite, and silica. For ultramafic materials, waters have a high concentration of Mg and are buffered to more alkaline pH.

Precipitatated minerals forming on the rock and/or in topographic lows from evaporation

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of transported fluids are Mg-carbonate bearing. The timing of this alteration relative to surface bodies of water is equivocal and may have been coeval with mineral deposition from surface waters in basins.

Examination of the deepest, and oldest stratigraphies on Mars leads to the conclusion, however, that alteration processes during the Noachian were profoundly influenced by impact cratering. Of the thousands of phyllosilicate exposures found on

Mars, the majority are associated with the ejecta, central peaks, and walls of craters spread over the southern highlands. The 600 m thick stack of Fe/Mg smectite-bearing breccias exposed by the Nili Fossae graben may typify the churning of Mars’ earliest crust by basin-sized impacts. The matrix is altered and exhibits evidence for fluid flow within (Mustard et al., 2009). Some of the breccia blocks are layered and preserve pre- existing, ordered stratigraphies of either lavas or sedimentary rocks. Elsewhere in the southern highlands, the “deep phyllosilicates”, ballistically ejected by smaller craters, show alteration under hydrothermal conditions (Chapter 5). Impact played substantial role in churning the earliest products of aqueous alteration, if not also in creating them via localized hydrothermal activity.

Outstanding questions

As VNIR images from OMEGA and CRISM continue to be collected and analyzed and new rovers and landers are sent to the surface, discoveries of still more minerals on Mars may yet be made. With the existing data, a number of questions remain outstanding.

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What is the abundance of alteration minerals?

Until rovers and landers visit some of the sites of mineralogic alteration discussed herein, orbital data remains the sole source for characterization of composition. As shown in Chapter 6, VNIR spectroscopy, applied carefully, is an accurate technique yielding few mineral misidentifications. But at present VNIR spectroscopy usually yields only 1-2 mineral identifications per pixel even though rocks are typically comprised of many more constituents. Thermal infrared spectroscopy, in contrast, can successfully model complex modal mineralogies (Ramsey & Christensen 2002) but does not reliably detect phyllosilicates on Mars. This is not simply an abundance issue; textural effects can successfully “hide” from TIR techniques even abundant minerals (Kirkland et al., 2003).

A goal of VNIR radiative transfer modeling is to move VNIR spectroscopy from the mostly qualitative analyses possible today—“this mineral is present”—towards quantitative assessment “minerals X, Y, and Z are present in mixture at x±a%, y±b%, and z±c%”. First estimates applying the Shkuratov radiative transfer model on VNIR spectral data from OMEGA provide estimates of phyllsilicate abundances ranging from 5-65%

(Poulet et al., 2008), but the degree of accuracy and precision possible for this technique is currently ill-understood. Placing bounds on uncertainty is a key goal of future laboratory radiative transfer modeling studies of synthentic mixtures and natural rock samples and is necessary to have confidence in the application of radiative transfer modeling techniques to remotely sensed data.

The benefits of successful quantitative modeling of modal mineralogy from VNIR data would be great. The proportion of altered to unaltered phases in a rock unit increases with degree of alteration and can serve as an indicator of the intensity and duration of

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water chemically interacting with primary minerals. The quantity of hydrated minerals in rock units is also important for understanding Mars’ overall water budget, i.e. constraining the amount of Mars’ water and atmospheric volatiles sequestered in the crustal rock reservoir.

Why the diversity?

There are hundreds of paleolakes and valley systems on Mars (Cabrol & Grin,

1999; Irwin et al., 2005; Fassett & Head, 2008). Only a few dozen have mineralogically interesting sedimentary deposits (Grant et al., 2008; Swayze et al., 2008; Ehlmann et al.,

2008a; Wray et al., 2009; Milliken et al., 2010). There are hundreds of thousands of impact craters on Mars, but many have no associated alteration minerals. Alteration minerals result from a complex combination of thermodynamic and kinetic factors such as pressure, temperature, pH, oxidation state (Eh), and activities of anions and cations.

Local groundwater availability and chemistry, bedrock, and climatic factors determine these as does duration of interaction with fluids. Past aqueous chemical environments are indicated (Figure 1) based on qualitative analysis of orbital data. However, the new, detailed mineralogic data also provides a wealth of information for geochemical modelers to use the discovered minerals and their assemblages to more tightly constrain geochemical conditions and the rates and processes of mineral formation on early Mars.

Efforts to reconcile the new mineralogic dat with models of alteration processes have begun for impact crater hydrothermal systems (e.g. Schwenzer & Kring, 2009), but progress needs to be made in reconciling previously promulgated models for lacustrine deposits (Catling, 1999), volcanic hydrothermal activity (Griffith and Shock, 1997;

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Zolotov, 2007), and near surface weathering (Chevrier et al., 2007) so as to establish more detailed constraints on the nature of Mars' earliest environments.

What was the role of impacts on early Mars and its habitats?

Unlike Earth, Mars preserves a substantial portion of its crust from the time of the time of heavy bombardment, and alone among the terrestrial planets, it also records a record of how a volatile-rich planet responds to impact. However, the age of the environments we are able to access from this orbital record is not entirely clear. Are most rocks on Mars from 3.7-3.9 Ga with previous records destroyed by an intensive heavy bombardment, or do more ancient surfaces persist? (e.g. Nimmo & Tanaka, 2005).

This timing aspect has important implications for understanding the nature of aqueous alteration and it evolution through time. For example, valley networks are well preserved from around the time of the Noachian-Hesperian boundary (~3.7 Gyr). But were they also active earlier and the record is simply destroyed (or preserved only in layered breccia blocks)?

There is a need for greater analog studies of neutral to alkaline weathering environments where Fe/Mg smectites such as those formed on Mars are found. While compelling from a habitability perspective, some findings appear not entirely consistent with exclusively near-surface phyllosilicate formation during the Noachian. For example phyllosilicates are rarely associated with the Noachian valley networks. These were sites of enhanced precipitation and runoff, and thus would be expected to be sites of enhanced weathering. Instead, impacts and lava flows may have been the prime drivers of phyllosilicate formation as subsurface waters interacted with rocks and impacts liberated volatiles (e.g. Segura et al., 2002) to interact with rocks.

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Which environments might preserve evidence for life?

The first billion years hosted multiple aqueous environments, differing in aqueous geochemistry and geologic setting. When sending landed missions to visit the Martian surface with biological goals—find carbon or other biosignatures—it is apparent there are multiple choices of sites. Scientifically, an intense debate rages over (a) what types of environments early microbial organisms might have been most likely to inhabit and (b) in which types of environments biosignatures of past life--organic carbon, morphologic fossils, mineral disequilibria, or distinctive isotopic fractions--are most likely to be preserved over three billion years. The answers to these questions weigh importantly in decisions concerning landing sites for sample return. Terrestrial examination of biomarkers has largely focused on near-surface shallow or deep water environments. The earliest fossil evidence for life on Earth is mostly photosynthetic, though any evidence for fossil life much older than ~2 Gyr is in dispute (Grotzinger & Knoll, 1999; Schopf,

2006). In contrast, sequencing the “tree of life” from 16s RNA indicates the oldest life forms were chemosynthetic , thermophillic organisms (Woese, 1987; DiGuilio, 2003;

Shimizu et al., 2007). The nature of preservation of modern day relatives to these organisms in hydrothermal and deep, buried systems needs to be understood before we can have confidence applying terrestrial metrics for biosignature preservation to ancient

Mars.

Why the global change?

Finally, while it is clear that geoscientists should think locally on Mars to understand the origins of particular aqueous environments and their habitability potential, evidence for the global nature of changes in aqueous processes through time has only

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become clearer. Why did this occur? Was Mars’ period of surface habitability brief, perhaps driven by release of volatiles by impacts (Segura, 2002) or release of sulfurous greenhouse gases by volcanism (Bullock and Moore, 2007; Johnson et al., 2008)? What role did internal processes such as cooling of the core, loss of the dynamo and increased solar stipping have on the loss of volatiles? In the next 5-10 years, with data from the next generation of atmospheric orbiters and revised estimates for the amount of water sequestered as minerals in Mars’ crust, we should arrive at new estimates for timing and reasons for the loss of Mars hydrologic cycle.

Implications for Mars exploration

A rover and two orbiters are planned for Mars from 2011-2016. Each landing site under consideration for the 2011 Mars Science Laboratory (MSL) has phyllosilicates, so this rover mission will provide the first opportunity to study rocks that date from Mars’ earliest epoch of alteration in-situ, at a location with clear geologic context. Following

MSL, two successive orbiters will contribute data toward understanding the processes, structure and composition of Mars’ atmosphere. In 2013, MAVEN will measure isotopic signatures, permitting quantification of volatile loss through time. The 2016 ExoMars orbiter will measure the abundance of trace gas species, including recently detected methane, to characterize any potential sources and sinks.

An aspect of the past 20 years of the Mars Exploration Program is the high degree of synergy between orbital and landed missions with sequential, complimentary investigations building off the findings of previous missions. From orbit, present mapping of the stratigraphy of mineralogic units would benefit greatly from higher-

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resolution stereo data. Future orbital mapping with higher resolution infrared imaging coupled with few meters-scale stereo and morphologic data would offer permit more refined tracking of changing environmental conditions. It remains the case, however, that rock fabrics and the textural relationships at the grain scale are critically important to deciphering past geologic processes and environmental histories. Though orbital technology continues to improve, landed missions, and perhaps eventually human missions, will continue to provide leaps forward due to their ability to examine records of geologic processes on the nm to m scale and with a far higher degree of certainty in compositional measurements.

Out of the data of the recent years, three key themes emerge for design of future landed missions: precision, mobility, and diversity. With orbital capabilities now permitting detailed traverse planning and the identification of outcrops of highest importance to a few meters-scale precision, the challenge for landing systems technology is to fully exploit this capability by permitting landings near the targets of highest interest. Extensive rock outcrops preserve the best record of paleocondtions; hence, what of is of keenest science interest is not most conducive to safe landings. At present, the smallest landing ellipse is 20 km in diameter. Under the most aggressive driving scenarios, this translates to at least 4 months of driving to exit an ellipse. Precision landing capabilities to within 5-10 km can cut drive times to ellipse exit to a matter of weeks and would allow access to identified regions of highest interest.

Mobility remains essential to move the science instrument package from the landing ellipse, selected for safety, to target regions, selected for their greatest science interest. Mobility is also critical for assessment of environmental change since this

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requires physical traverses of stratigraphic boundaries to examine changing mineralogy, chemistry, and morphology of rock units. Except for dedicated monitoring for seismic and weather studies, little would be gained by future immobile landers on Mars, and if such stationary platforms are placed, efforts should be made to couple with landings of small rovers.

Finally, and perhaps most profoundly, findings of the last 5 years point to the incredible diversity of aqueous environments on ancient Mars. Nearly a dozen distinctive settings of aqueous processes existed on ancient Mars, and ice-rich regions of modern near-polar Mars are also of interest. Based on our terrestrial experience, each of these has the potential to host a habitable environment. Our understanding of the climatic and geologic processes that generated any of these environments is as of yet incomplete, as is our understanding of which are most likely to preserve biosignatures through time. The benefits are sample return are substantial: dating of samples, application of laboratory sample prep capabilities, microanalytical capabilities and detailed isotopic studies.

However, given both the diversity of ancient Mars and the lessons learned from difficult, often equivocal analyses of samples from ancient Earth, one must be cautious about what a single set of samples is likely to tell us about Mars geochemistry and the potential for life on Mars. Many mid-class rovers and their in-situ analyses rather than a single sample return may in fact provide a more comprehensive picture of ancient Mars environments.

One must balance the data gained from each mission scenario with their costs.

Final thoughts

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Mars has a rich record from its first billion years available for exploration. By combining orbital spectroscopic data with in-situ studies from the next generation of rovers, we can make substantial progress in understanding the history of water on Mars, its evolution over billions of years, and perhaps, whether Mars ever hosted the geochemical environments necessary to sustain life. Mars’ intact rock record pushes backward our ability to access early terrestrial planetary environments at the time of life’s origins, allowing us to understand its beginnings and, perhaps, its future. Such work brings us closer to understanding the early environments of our own Earth and, indeed, what is required to sustain a habitable planet through time.

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Table 1. Minerals resulting from aqueous alteration, discovered on Mars through January 2010. All mineral phases are observed from orbit with VNIR spectroscopy except for phases in italics, which have only been reported by other orbiting and landed missions.

Class Group/mineral Formula References

Fe/Mg smectites (e.g., Poulet et al., 2005; (Ca, Na) (Fe,Mg, Al) (Si, Al) O (OH) nontronite, saponite) 0.3-0.5 2-3 4 10 2 Mustard et al., 2008

Poulet et al., 2005; montmorillonite (Na,Ca) (Al,Mg) (Si O )(OH) 0.33 2 4 10 2 Mustard et al., 2008

kaolin group minerals Bishop et al., 2008; Al Si O (OH) (e.g. kaolinite, halloysite) 2 2 5 4 Ehlmann et al., 2009 Phyllosilicates Mustard et al., 2008; chlorite (Mg,Fe2+) Al(Si Al)O (OH) 5 3 10 8 Ehlmann et al., 2009

serpentine (Mg, Fe)3Si2O5(OH)4 Ehlmann et al., 2009

high charge Al,K Mustard et al., 2008; phyllosilicate (e.g. KAl AlSi O (OH) 2 3 10 2 Ehlmann et al., 2009 muscovite or illite)

Clark et al., 2008; prehnite Ca Al(AlSi O )(OH) 2 3 10 2 Ehlmann et al., 2009 Other

hydrated analcime NaAlSi2O6·H2O Ehlmann et al., 2009 silicates Milliken et al., 2008; opaline silica SiO ·H O 2 2 Ehlmann et al., 2009

magnesium carbonate MgCO Ehlmann et al., 2008b (e.g. magnesite) 3 Carbonates Phoenix lander (Boynton Calcium carbonate CaCO 3 et al., 2009)

Fe/Mg mono- and poly- (Fe,Mg)SO · nH O Gendrin et al., 2005 hydrated sulfates 4 2

gypsum CaSO4· 2H2O Langevin et al., 2005

Sulfates alunite KAl3(SO4)2(OH)6 Swayze et al., 2008

3+ jarosite KFe 3(OH)6(SO4)2 Milliken et al., 2008

Swayze et al., pers. not a named mineral FeSO (OH) 4 comm.

TES (Osterloo et al., Chlorides chlorides (metal)Cl 2008)

Phoenix Lander (Hecht Perchlorates perchlorates (Mg,Ca)(ClO ) 4 2 et al., 2009)

Christensen et al., 2000; hematite Fe O 2 3 Bibring et al., 2007 Fe Oxides goethite FeO(OH) Farrand et al., 2009

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Figure 1. A timeline of the chemical environments of ancient Mars, adapted from Murchie et al. (2009) with mineralogic epochs from Bibring et al. (2006)

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