I Identification and Characterization of Martian Acid-Sulfate Hydrothermal

Home , Barth (crater), Bernard (crater), Bounce Rock, Gill (Martian crater), Maraldi (Martian crater), Matijevic Hill

Identification and Characterization of Martian Acid-Sulfate Hydrothermal Alteration: An

Investigation of Instrumentation Techniques and Geochemical Processes Through

Laboratory Experiments and Terrestrial Analog Studies

Sarah Rose Black

B.A., State University of New York at Buffalo, 2004

M.S., State University of New York at Buffalo, 2006

A thesis submitted to the

Faculty of the Graduate School of the

University of Colorado in partial fulfillment

of the requirement for the degree of

Doctor of Philosophy

Department of Geological Sciences

2018

i This thesis entitled: Identification and Characterization of Martian Acid-Sulfate Hydrothermal Alteration: An Investigation of Instrumentation Techniques and Geochemical Processes Through Laboratory Experiments and Terrestrial Analog Studies written by Sarah Rose Black has been approved for the Department of Geological Sciences

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Dr. Brian M. Hynek

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Dr. Alexis Templeton

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Dr. Stephen Mojzsis

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Dr. Thomas McCollom

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Dr. Raina Gough

Date: ______

The final copy of this thesis has been examined by the signatories, and we find that both the content and the form meet acceptable presentation standards of scholarly work in the above mentioned discipline.

ii Black, Sarah Rose (Ph.D., Geological Sciences)

Identification and Characterization of Martian Acid-Sulfate Hydrothermal Alteration: An

Investigation of Instrumentation Techniques and Geochemical Processes Through

Laboratory Experiments and Terrestrial Analog Studies

Thesis directed by Associate Professor Brian M. Hynek

Abstract

Hydrothermal acid-sulfate alteration is a common process in volcanic systems on Earth, and it may be inferred that this process played an important role in Mars’ geologic history as well.

Several areas have been identified on Mars with minerals which are characteristic of acid-sulfate alteration: hydrated silica, sulfates, phyllosilicates, and Fe-oxides. Relic hydrothermal systems will play a key role in future investigations of Mars and the search for biosignatures. It is necessary to develop a detailed understanding of the mineral assemblages which form in these environment, and the geochemical processes by which they arise. This dissertation work addresses our ability to confidently and thoroughly characterize hydrothermal acid-sulfate mineral assemblages using rover-deployed instrumentation methods, and the role of high Fe parent basalts on secondary mineralogy.

Analysis of hydrothermal deposits was completed using Mars analog instrumentation to identify strengths and weaknesses for each method. VNIR, XRD, and Raman laser spectrometers analyzed 100 hydrothermally altered terrestrial samples. Results indicate phyllosilicates may be detected in XRD without any additional sample preparation when present at ≥ 10 wt %. VNIR was particularly useful for the identification of phyllosilicates and silica; however, these deposits are easily missed when observed with limited rover VNIR. For more robust phyllosilicate detection

iii and “ground truthing” of orbital data, VNIR spectrometers on future rover payloads should cover the entire 300 – 2500 nm range.

Previous work has shown that parent rock composition plays a significant role in secondary mineral assemblage, and Icelandic basalts are among the closest terrestrial analogs for Mars. Our examination of the role of primary Fe content on secondary mineralogy found ~60 % Fe natroalunite forming in Iceland, adding to the growing list of natural systems containing intermediate alunite group minerals. There is a direct correlation between parent Fe and the abundance of Fe-bearing secondary minerals. Projected trends from our laboratory investigations indicate 20 – 32 wt % Fe-bearing mineralogy should form in Martian systems, similar to products from our terrestrial basalts. Alteration of crystalline “Mars” basalt produced both Fe-sulfates and

Fe-oxides while all others contained only Fe-sulfates, indicating a transition to co-precipitation at

> 17.0 wt % parent FeOT.

iv Acknowledgements

Many thanks to Brian Hynek for his support and making this project possible, as well as the current and former members of the mighty SPACECATS lab, fellow LASP-ians, and friends for their never ending support, encouragement, and laughs: LG Beckerman, John Gemperline,

Rachael Hoover, Kara Brugman, Kaitlyn Garifi, Katie Rempfert, Stephanie Junior, Meredith

Hazelton, Molly DiCroce, Zarah Brown, Morgan Rehnberg, Marek Slipski, Courtney Peck, Emma

Marcucci, Mike Lotto, Will Nelson, Becca Thomas, and Ramy El-Maarry. Numerous people have contributed to my knowledge and growth as a scientist throughout this program, to whom I am extremely grateful: my committee (Alexis Templeton, Tom McCollom, Steve Mojzsis, and Raina

Gough), Kathy Kierian-Young, Aileen Yingst and the members of the GHOST team, Matt

Chojnacki, Tracy Gregg, Graham Lau, Chris Donaldson, Eric Ellison, Julien Allaz, Jess Larsen, and the truly outstanding CU Boulder geology faculty. And of course, nothing would have been possible without my incredible family: Kathy, Steve, Julie, Dan, and Ryan, and my truly wonderful partner, Jake Aho. Thank you all. I could not have done it without you.

Chapters 2 – 4 have been published, submitted, or will soon be submitted in peer-reviewed scientific journals. The acknowledgements for each section are as follows:

Chapter 2: This work was funded by a Geological Society of America Graduate Student

Research Grant, the Lewis and Clark Fund for Exploration and Field Research in Astrobiology, a

NASA Early Career Fellowship to B. M. Hynek, and NASA awards NNX14AN36G and

NNX14AG90G. Many thanks to Rebecca Thomas for providing valuable feedback, Kathy

Kierian-Young, Emma Marcucci, Kara Brugman, Thomas McCollom, Eric Ellison, LG

Beckerman, Rachael Hoover, Stephanie Junior, and Jordan Ludyan for their assistance with

v gathering and processing data, and Guillermo Alvarado, and Geoffroy Avard for their assistance in the field.

Chapter 3: This work was supported by a Geological Society of America Graduate Student

Grant; the American Philosophical Society Lewis and Clark Fund for Exploration and Field

Research in Astrobiology, and NASA Habitable Worlds grant NNX15AP15G to LJM and BMH.

Many thanks to Jordan Ludyan for assistance with XRF analysis at UWM.

Chapter 4: This work was supported by a Geological Society of America Graduate Student

Grant; the American Philosophical Society Lewis and Clark Fund for Exploration and Field

Research in Astrobiology, and NASA Habitable Worlds grant NNX15AP15G to BMH. Many thanks to Jordan Ludyan for assistance with XRF analysis of parent rocks at UWM.

vi Table of Contents

1. Introduction …………………………………………………………………..…………….. 1 1.1. Observation of Mars: A History …………………….……………………..……….….. 2 1.1.1. Telescopic observations ………………………………………………………..… 2 1.1.2. Orbital and in situ spacecraft observations …………………………...………..… 3 1.1.3. Upcoming missions ………………………………………………..…………..… 6 1.2. The geology of Mars ……………………………..……………………...…...………..… 7 1.2.1. A brief geologic history …………..………………………………..…………..… 7 1.2.2. Volcanism and iron-rich basalts ……………….……………………….………… 9 1.2.3. Evidence for surface and subsurface water ………………………..…….……… 11 1.2.3.1. Geomorphological evidence ………………………………………..…….... 11 1.2.3.1.1. Valley networks and deltas ………………………………………..…. 11 1.2.3.1.2. Outflow channels ……………………………………………….....… 12 1.2.3.1.3. Gullies ……………………………………………………………..… 13 1.2.3.2. Mineralogical evidence ………………………………………………….… 13 1.2.3.2.1. Phyllosilicates ……………………………………………………..… 14 1.2.3.2.2. Sulfates …………………………………………………………..…... 15 1.2.3.2.3. Fe-oxides/hydroxides ……………………………………………..…. 16 1.2.3.2.4. Hydrated silica ……………………………………………………..... 17 1.3. Hydrothermal systems on Mars …………..……………………….……..…………… 17 1.3.1. Home Plate, Gusev crater …………………………………………………...... … 18 1.3.2. Eastern Coprates Chasma, Valles Marineris ……………………………..…..… 19 1.3.3. Noctis Labyrinthus, Valles Marineris ……………………………….....……..… 19 1.3.4. Cross crater, Terra Sirenum …………………………………………..…..…..… 19 1.3.5. Nili Patera, Syrtis Major ……………………………………………….....…..… 20 1.4. Acid-sulfate alteration: Processes and products ……..……………………...……..… 20 1.5. Dissertation research ………………………………….…………………..…….…..… 22

2. Characterization of terrestrial hydrothermal alteration products with Mars analog instrumentation: Implications for current and future rover investigations …..…..…...... 25 2.1. Introduction ………………………………………………………………….…...…… 25 2.1.1. Hydrothermal alteration on Mars ……………………………………….…..…… 25 2.1.2. Instrumentation ……………………………………………………………...... … 28 2.1.2.1. Visible Near-Infrared (VNIR) and Short Wave Infrared (SWIR) Reflectance Spectroscopy ……………………………………………………………………..… 28 2.1.2.2. X-Ray Diffraction (XRD) ………………………………………………..... 32 2.1.2.3. Raman Laser Spectroscopy ………………………………….…………..… 33 2.1.3. Characterization of terrestrial hydrothermal deposits …………………..……...… 35 2.1.4. Field sites ………………………………………………………………...……… 38 2.1.4.1. Poás Volcano, Costa Rica ……………………………………………..…... 40 2.1.4.2. Turrialba Volcano, Costa Rica ………………………………………..…… 41 2.1.4.3. Cerro Negro Volcano, Nicaragua ……………………………………..…… 41 2.1.4.4. Momotombo Volcano, Nicaragua ……………………………….…..…….. 42 2.1.4.5. Telica Volcano, Nicaragua ……………………………………….…..……. 42 2.1.4.6. Krafla Volcano, Iceland ……………………………………………..…….. 43

vii 2.1.4.7. Landmannalaugar Volcano, Iceland …………………………….……..…... 43 2.2. Methods …………………………………….……………………………………..…… 44 2.2.1. Sampling …………………………………………………………….………..…. 44 2.2.2. Analysis ……………………………………………………………..………..….. 44 2.2.2.1. VSWIR …………………………………………………………….…..…... 45 2.2.2.2. XRD ……………………………………………………………………..… 46 2.2.2.3. Raman …………………………………………………………………..…. 47 2.3. Results ……………………………………………..………………………………..….. 48 2.3.1. SiO2 …………………………………………………………………………..….. 51 2.3.2. Elemental Sulfur ………………………………………………………………..... 54 2.3.3. Sulfates ………………………………………………………………………..…. 55 2.3.4. Oxides and Hydroxides ………………………………………………………..… 57 2.3.5. Zeolites ……………………………………………………………………..……. 59 2.3.6. Phyllosilicates ………………………………………………………………..….. 59 2.4. Discussion ……………………………………….…………………………………..…. 62 2.4.1. VSWIR …………………………………………………………………………... 62 2.4.2. XRD …………………………………………………………………………..…. 63 2.4.3. Raman ………………………………………………………………………..….. 64 2.4.4. Implications for Mars ………………………………………………………..…... 66 2.4.4.1. MER Spirit at Gusev crater ……………………………………………..…. 66 2.4.4.2. MER opportunity at the Matijevic formation, Endeavour crater ……..…… 73 2.4.4.3. MSL Curiosity at Gale crater …………………………………………..….. 74 2.4.4.4. ExoMars, Mars 2020, and beyond ……………………………………..….. 76 2.5. Conclusions ………………………………………….…………………………..…….. 79

3. Bulk mineralogy of surficial hydrothermal acid-sulfate deposits at Námafjall, Þeistareykir Geothermal Field, and Hengill Volcano, Iceland: Implications for the identification and interpretation of hydrothermal deposits on Mars ……………...……….. 82 3.1. Introduction ……………………………………………….…………..…………..…... 82 3.1.1. Hydrothermal systems on Mars ………………………………...... …… 82 3.1.2. Terrestrial studies of hydrothermal systems …………………………...….…..…. 83 3.1.3. Geologic setting of sample sites ………………………………………..………... 86 3.1.3.1. Námafjall geothermal field, Krafla Region ……………………..…..…..…. 87 3.1.3.2. Þeistareykir geothermal field, NVZ …………………………..……...…..… 88 3.1.3.3. Nesjavallavir Power Plant, Hengill Volcanic Complex …………..…….….. 90 3.2. Methods ………………………………………………….………………..…..……..… 92 3.2.1. Sample collection ………………………………………………..………..…..…. 92 3.2.2. Analysis …………………………………………………………..……..….…..... 92 3.2.2.1. Fluid chemistry ………………………………………………..…....……… 92 3.2.2.2. Visible Near-Infrared (VNIR) Spectroscopy ………………..…….....…….. 93 3.2.2.3. X-Ray Diffraction (XRD) ………………………………..…………...... … 93 3.2.2.4. X-Ray Fluorescence (XRF) ………………………………...………..…..… 94 3.2.2.5. Electron Microprobe Wavelength-Dispersive Spectrometer (EPMA-WDS) 94 3.3. Results ……………………………………………………..…………………....…..….. 95 3.3.1. Primary basalt composition ………………………………………..………..…… 95 3.3.1.1. Námafjall geothermal field, Krafla Volcano …………..………….……..… 95

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3.3.1.2. Þeistareykir geothermal field, NVZ ………………………………..….…... 95 3.3.1.3. Nesjavallavir Power Plant, Hengill Volcanic Complex …………..……..… 96 3.3.2. Secondary mineralogy …………………………………………………..……..… 97 3.3.2.1. Námafjall geothermal field, Krafla Volcano ……………………...…..…… 97 3.3.2.2. Þeistareykir geothermal field, NVZ …………………………..……….…. 100 3.3.2.3. Nesjavallavir Power Plant, Hengill Volcanic Complex ……….…….…… 104 3.4. Discussion ………………………………………………………..………………….... 107 3.4.1. Secondary mineralogy …………………………………………………..….….. 107 3.4.1.1. Námafjall geothermal field, Krafla Volcano ……………………..…..…... 107 3.4.1.2. Þeistareykir geothermal field, NVZ ……………………………..……..… 110 3.4.1.3. Nesjavallavir power plant, Hengill Volcanic Complex ………...…..…….. 114 3.4.2. Implications for Mars hydrothermal deposits ………………………...…..…….. 119 3.4.2.1. Comparison to Mars mineralogy ………………………………..…..……. 119 3.4.2.2. Fe-rich natroalunite detections ………………………………..……..…… 120 3.5. Conclusions ……………………………………………………………..…...……….. 123

4. Hydrothermal acid-sulfate alteration of a synthetic Mars composition basalt: Implications for hydrothermal deposits on Mars …………………………………….....….. 125 4.1. Introduction ………………………………………………………………..……...…. 125 4.1.1. Hydrothermal systems on Mars …………………………………………..…..… 125 4.1.2. Martian basalt compositions ………………………………………..………..…. 126 4.1.3. Previous work …………………………………………….……..…………..….. 128 4.2. Methods ……………………………………………………………..…………...…… 132 4.2.1. Samples ………………………………………………………………..…..….... 132 4.2.1.1. Synthetic Mars Basalt (SMB) ………………………………..………...…. 132 4.2.1.2. Terrestrial parent materials ………………………………..…………..….. 134 4.2.2. Experimental setup ……………………………………………..…………..…... 136 4.2.3. Sample analysis ………………………………………………..…………..…… 137 4.2.3.1. Fluid chemistry ………………………………………..……………..…… 137 4.2.3.2. X-Ray Diffraction (XRD) …………………………..……………….....… 138 4.2.3.3. Visible Near-Infrared (VNIR) Spectroscopy …………………..……..….. 138 4.2.3.4. Scanning Electron Microscope Energy-Dispersive Spectroscopy (SEM-EDS) ………………………………………………………………… 139 4.3. Results …………………………………………………………………………..…….. 139 4.3.1. Closed system …………………………………………………………..…..…... 139 4.3.1.1. Crystalline synthetic Mars basalt (SMBxl) ……………………..……..….. 139 4.3.1.2. Glassy synthetic Mars basalt (SMBgl) ………………………..……..…… 142 4.3.1.3. High Fe terrestrial basalt (HN-BAS) ……………………….………..…... 143 4.3.1.4. Low Fe terrestrial basalt (CN-BAS) ………………………………..…...…143 4.3.1.5. Terrestrial basaltic andesite (B-AND) ……………………………..…..…. 144 4.3.1.6. Terrestrial andesite (AND) ………………………………………..…..….. 144 4.3.1.7. Terrestrial rhyolite (RHY) and obsidian (OBS) ………………..……...….. 144 4.3.2. Flow-through system …………………………………………………..….……. 145 4.3.2.1. Crystalline synthetic Mars basalt (SMBxl) ……………………..…..…….. 145 4.3.2.2. Glassy synthetic Mars basalt (SMBgl) ………………………..……..…… 149 4.4. Discussion ………………………………………………………………..…………… 152

4.4.1. Effect of parent lithology on secondary mineralogy ……………….……..……. 152 4.4.2. Evolution of open-system SMB alteration …………………………...…..……... 155 4.4.3. Application to Martian hydrothermal systems ………………………….………. 159 4.4.4. Implications for terrestrial analog studies ……………………………..……..…. 161 4.5. Conclusions ………………………………………………………………..……...….. 162

5. Discussion and Conclusions …………………………………………………...…..….… 164

References ………………………………………………………………………..………..…. 171

List of Tables

Table 1: Iron abundances in unweathered Martian basalts and common terrestrial analogs ...... 10

Table 2: Common hydrothermal alteration products ...... 21

Table 4: Terrestrial and Mars surface compositions of primary basalts, basaltic andesites, and rhyolites represented as oxide weight percents...... 40

Table 5: Abundance of mineral subgroups out of 100 Mars analog hydrothermal samples analyzed with VSWIR, XRD, and Raman...... 49

Table 6: Characteristic features used for identification with VNIR, XRD, and Raman...... 50

Table 7: Relative identification accuracy of each instrument for each major mineral class. Stated percent is out of the total number of times a mineral was identified in a sample (eg: if a mineral was identified by either VSWIR, XRD, and/or Raman in 50 of the 100 samples, and was identified via VNIR 25 of those 50 times, % = 25/50, or 50%; max. possible = 100%). In many cases, one or two instruments identified a mineral, while the others did not. In the event that all three instruments failed to identify a mineral phase, it is not considered for that sample...... 51

Table 8: Previous analog studies of Icelandic hydrothermal acid-sulfate systems ...... 85

Table 9: Elemental weight percent oxide compositions measured with XRF for fresh basalts collected in this study...... 95

Table 10: Secondary alteration mineralogy identified with VNIR and XRD at Námafjall geothermal field ...... 98

Table 11: Fluid geochemistry: measured ion concentrations and saturation states for common hydrothermal minerals, and field ORP values ...... 100

Table 12: Secondary alteration mineralogy identified with VNIR and XRD at Þeistareykir geothermal field ...... 102

Table 13: Secondary alteration mineralogy identified with VNIR and XRD at Nesjavallavir power plant, Hengill...... 105

Table 14: Elemental compositions of unweathered Martian basalts and commonly-utilized terrestrial analogs ...... 128

Table 15: Previous laboratory-based acidic alteration studies ...... 131

Table 16: Elemental compositions of parent materials measured by XRF analysis ...... 132

Table 17: Mineral components of parent rocks as determined by XRD, EPMA, and/or petrographic thin section ...... 133

Table 18: Experimental conditions for each laboratory set up ...... 137

Table 19: Closed-system secondary mineralogy as seen in VNIR, XRD, and SEM-EDS ...... 141

Table 20: Mineral phases with a calculated SI > -0.4 in closed-system alteration fluids ...... 142

Table 21: Flow through SMBxl secondary mineralogy as seen in XRD and SEM-EDS ...... 147

Table 22: Mineral phases with a calculated SI > -0.4 in flow through SMBxl alteration fluids 149

Table 23: Flow through SMBgl secondary mineralogy as seen in XRD and SEM-EDS ...... 151

Table 24: Mineral phases with a calculated SI > -0.4 in flow through SMBgl alteration fluids 152

Table 25: Summary of secondary mineral phases in all experimental setups. Bold font indicates Fe-bearing secondary minerals ...... 153

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List of Figures

Figure 1: Timeline of the geologic activity on Mars from Carr and Head (2010) …………….… 8

Figure 2: Geomorphological evidence for flowing surface water; Left to right: valley networks, outflow channels, deltas, and gullies …………………………………………………..…. 11

Figure 3: Global distribution of aqueous secondary mineralogy across Mars, from Ehlmann et al. (2014) ……………………………………………………………………………………... 14

Figure 4: Sampling locations in Costa Rica (Poás, Turrialba), Nicaragua (Momotombo, Telica, Cerro Negro), and Iceland (Krafla, Landmannalaugar). Stars represent capital cities …... 39

Figure 5: Primary lithology compositions of individual samples collected from locations used in this study (Poás, Turrialba, Cerro Negro, Momotombo, and Telica) and previously reported compositions from Landmannalaugar and Krafla, compared to measured Martian surface compositions. Parent rock compositions from these sampling locations are similar to those encountered at various locations in situ on Mars, from orbital measurements, or from SNC meteorites. Gale clast data from Sautter et al. (2015). Cerro Negro, Momotombo, and Telica data from Hynek et al. (2013). Krafla data from Nicholson and Latin (1992). Landmannalaugar data from Blake (1984). Poás and Turrialba compositions from this study. Figure adapted from McSween et al. (2009) …………………………………………..…. 39

Figure 6: A) The TerraSpec 4 High Resolution Reflectance Spectrometer deployed in the summit crater of Poás Volcano, Costa Rica; B) The Terra X-Ray Diffraction Spectrometer deployed in the summit crater of Turrialba Volcano, Costa Rica; C) The Horiba LabRAM HR Evolution Raman Microscope-Spectrometer in the CU Boulder Laboratory for Environmental and Geological Studies ………………………………………………………………………... 45

Figure 7: Detail of SiO2 identification in VSWIR, XRD, and Raman using representative examples. The same samples have not been used for each method. Rather, the best possible example is shown. Quartz, cristobalite, and tridymite are relatively featureless in the VNIR range and need Mid-IR or thermal emission spectra for identification (Clark et al., 2013; Michalski et al., 2003; PDS Geosciences Node and CRISM Spectral Library Working Group, 2014). Opal, however, exhibits distinct H2O/OH and Si-OH absorption bands in the VNIR range (panels A and B). In panel B, VSWIR spectra are displayed as continuum-removed for more accurate band center locations. Red VSWIR spectra are reference spectra from USGS and CRISM libraries. Quartz, cristobalite, tridymite, and amorphous hydrated SiO2 were all identified with XRD (panel C), and opal, cristobalite, and quartz were identified with Raman. It is possible to identify tridymite with Raman (Lafuente et al., 2015); however, no Raman detections occurred in this study ………………………………………………………….. 53

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splibo6 library spectra for K-alunite, and pink is the McCollom et al., (2014) Fe-rich (10%) natroalunite library spectra. VSWIR spectra a and b (CR13 Turrialba 3 and I13 P0 yl) display the characteristic sulfur shoulder, while spectra c (CR13 Poás Hot 5) does not, despite containing ~ 19 wt% sulfur via XRD analysis ……………………………………………. 54

Figure 9: Details of sulfate identifications in VSWIR, XRD, and Raman using representative samples – easily identifiable using all three methods. Highlighted areas in panel A are shown in more detail in panels B, C, and D. In panels B, C, and D, VSWIR spectra are displayed as continuum-removed for more accurate band center locations ……………………………. 56

Figure 10: Details of oxide/hydroxide identifications in VSWIR, XRD, and Raman using representative samples. Highlighted area in panel A is shown in more detail in panels B and C. In panel C, VSWIR spectra are displayed as continuum-removed for more accurate band center locations. The Fe-oxides/hydroxides hematite and goethite are easily identifiable with all three methods. The Ti-oxide anatase not identifiable with VNIR due to the lack of features in the 350 – 2500 nm range ……………………………………………………………..… 58

Figure 11: Details of phyllosilicate identifications in VSWIR and XRD using representative samples. Highlighted areas in panel A are shown in more detail in panels B and C. Phyllosilicate identification in VSWIR (panels A-C) via the ~1.4 m OH/H2O overtone absorptions (panel B), and metal-OH combination from 2.1 to 2.4 m (panel C). In panels B and C, VSWIR spectra are displayed as continuum-removed for more accurate band center locations. In sinter containing rocks, the ~1.4 m band can be masked by the siliceous sinter signal and is not always a reliable diagnostic region. 2.1 to 2.4 m Metal-OH bands provide the most useful diagnostic tool in the VSWIR range. XRD identification of phyllosilicates is often limited to small peaks in the 5 – 15 2 range (panel D), but phyllosilicate-rich samples such as I13 KL cs min2 may display secondary peaks as well. Phyllosilicate identification is not possible with a 532 nm Raman due to interference from laser induced fluorescence ... 61

Figure 12: Laboratory spectra of phyllosilicate-bearing samples collected for this study, manually deconvolved to MER/MSL PanCam bands (panel A). All of the collected samples shown here contain kaolinite, as evidenced by the strong ~2100/2200 nm doublet in panel C (continuum-removed spectra for more accurate location of band centers). Nontronite (I13 KL red and I13 Nama U rd 1) and potentially saponite (CR13 Poás Lake 3) are also present in various samples, and may be identified by the ~2286 nm (nontronite) and ~2315 nm (saponite) absorption bands, also shown in panel C. Using MER/MSL PanCam band locations (panel A), kaolinite is not identifiable in any of these samples due to the lack of characteristic absorption bands at shorter wavelengths, and nontronite would be incorrectly identified as being present in both the CR13 Poás Lake 3 and CR13 Poás Lake 4 samples. At shorter wavelengths, the VNIR spectra appear to be dominated by Fe-oxides – the I13 KL red and I13 Nama U rd 1 samples both display the 532 and 834 nm Fe electronic transition absorptions indicative of hematite (panels A and B), which mask the 934 nm absorption feature that may indicate the presence of nontronite, but may also be due to other minerals such as Fe-bearing montmorillonite, or the Fe-hydroxide, lepidocrocite ………………… 69

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Figure 13: Laboratory spectra of amorphous SiO2-bearing samples collected for this study, deconvolved to MER/MSL PanCam bands (panel A). All of the collected samples shown here contain opal or siliceous sinters, as evidenced by the strong ~1410/1460 nm doublet in panel C, asymmetrical ~1900 nm band in panel D, and ~2210/2260 nm doublet in panel E (continuum-removed spectra for more accurate location of band centers). Presence of amorphous SiO2 is also confirmed via XRD for each sample, and distinctions between opal- A, opal-CT, and sinters were made using VNIR library spectra as well as 1410, 1460, 1900, and 2200 nm band depth and ratios as described by Rice et al. (2013). Although amorphous SiO2 is present, it is often not detectable with MER/MSL PanCam bands – as evidenced by the lack of downturn at 1009 nm. Often this lack of a 1009 nm absorption is due to the presence of hematite, goethite, or other Fe-bearing materials, which dominate the VNIR spectra at shorter wavelengths, as seen in samples CR13 Poás Fumarole 12 and I13 Nama U or1, I13 Nama Core1, CR13 Poás Fumarole 11, and CR13 Turritop 1. Of this subset, the only samples that would appear to contain amorphous SiO2 at MER/MSL resolution are CR13 Turritop1 and CR13 Turrialba N7. Hydrated SiO2 in all other samples would likely not be identified ………………………………………………………………………………..… 72

Figure 14: Top: Sampling locations across Iceland; bottom: Fe, Mg, Al, Na, and K oxide abundances measured by XRF and normalized to SiO2 for unaltered Martian basalts in Gusev (H. Y. McSween et al., 2006; Ming et al., 2008) and Meridiani (Rieder et al., 2004), unaltered Shergottite basalts (Hurowitz et al., 2005; Lodders, 1998; Rubin et al., 2000), commonly- utilized terrestrial analog sites in Hawaii (Chemtob and Rossman, 2014; Morris et al., 2000), Nicaragua (Hynek et al., 2013), and Costa Rica (Black and Hynek, 2017), and Icelandic sampling locations for this study …………………………………………………………. 86

Figure 15: Sampling locations (red stars represent mudpots, and green diamonds indicate fumaroles) at Námafjall. Image courtesy of Google Earth ………………………………. 88

Figure 16: Sampling locations at the Þeistareykir Geothermal Field. Image courtesy of Google Earth …………………………………………………………………………………….… 90

Figure 17: Sampling locations (white outlined areas in insets A and B) at the Nesjavallavir Power Plant on the flanks of the Hengill Volcanic Complex. Satellite images courtesy of Google Earth ………………………………………………………………………………………. 91

Figure 18: Representative samples from Námafjall with annotated XRD patterns and VNIR reflectance spectra; Far right: cm-scale dried hand samples; Al-phy – Al-phyllosilicate; Al- sulf – Al-sulfate; SiO2 – amorphous SiO2; An – anhydrite; F – primary feldspar; Fe – iron; 61 – ~60% Fe natrojarosite; Gy – gypsum; Hal – halotrichite group; H – hematite; K – kaolinite; Mo – montmorillonite; Na – natroalunite; Non – nontronite; P – primary pyroxene; Pk – pickeringite; R – rhomboclase; Roz – rozenite; Smec – smectite; S – sulfur; V – voltaite ………………………………………………………………………………... 99

G – goethite; Gy – gypsum; Hal – halotrichite group; H – hematite; J – jarosite; K – kaolinite; Mo – montmorillonite; Pk – pickeringite; Py – pyrite; R – rhomboclase; S – sulfur; Vr – vermiculite …………………………………………………………………………….… 103

Figure 20: Representative samples from Hengill with annotated XRD patterns and VNIR reflectance spectra; Far right: cm-scale dried hand samples; A – primary augite; Al-phy – Al- phyllosilicate; At – anatase; SiO2 – amorphous SiO2; Ac – analcime; D – primary diopside; Fe – iron; G – goethite; Gy – gypsum; Hal – halotrichite group; J – jarosite; K – kaolinite; Mo – montmorillonite; Q – quartz; S – sulfur …………………………………………... 106

Figure 21: Mineralogy of Námafjall gradient samples BMH N3 (left) and Nama U8 (right); stars indicate vent location ……………………………………………………………………. 110

Figure 22: Alteration environments at the Þeistareykir mudpots. Mudpots 1 and 2 were observed to have similar pH (pH 3), temperature (80-95C), and fluid:rock ratios, while mudpot 3 had a lower fluid:rock ratio, pH of 1.5, and measured temperatures of 60-98C ……………. 111

Figure 23: Coarse-grained gray (specular) hematite at Þeistareykir mudpot 3 ………………. 113

Figure 24: Multicolor stream system at Hengill-2. Letters indicate the location of close up images. Unless otherwise noted, scale bars are ~1 meter. A) Red stream; B) Gray stream; C) White stream; D) Confluence of gray and white streams; E) Confluence of red and gray/white streams; F) A strong oxidation front is visible with abundant pyrite located just below a surface crust on the side of the red stream (location of image F is noted in inset A) ….... 116

Figure 25: Streambed mineralogy at Hengill-2 ………………………………………………. 117

Figure 26: Eh/pH diagram of sulfur speciation (constructed using Geochemists Workbench Act2 program) in the white stream at 22C, pH 5, using measured fluid composition from the white stream source. Spring water emanates from the subsurface with reducing conditions at point A, and progresses along the equilibrium path to point B (observed Eh at white spring source, ORP = 0.217 V). S0 is formed (yellow region) as the solution moves along this geochemical pathway (black circle). Blue fields indicate aqueous species …………………………... 117

Figure 27: Fe-rich (~60% Fe in the B site) (natro)alunite in the Nama U4 sample, as seen in XRD and Raman. Gypsum XRD pattern from RRUFF (Lafuente et al., 2015). Natroalunite, natrojarosite, and intermediate XRD and Raman library spectra/diffraction patterns from McCollom et al. (2014). SEM image shows areas (white circles) with EDS spectra consistent with Fe-bearing alunite group minerals. The EDS spectra for point 1 has a roughly 1.5:1 Al:Fe ratio, while point 2 has closer to 1:1 Al:Fe …………………………………………….... 123

Figure 28: SMB & terrestrial parent materials in thin section and SEM; ol = olivine, px = pyroxene, pl = plagioclase feldspar, qtz = quartz, b = biotite, c = chlorite; Top: Backscattered electron (BSE) images of parent basalts; Bottom left: Crystalline parent rocks (SMB and terrestrial) in thin section, FOV = 2000 µm across; Bottom right: Average major mineral composition of

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parent basalts, measured by EPMA-WDS and averaged over 10 – 15 unaltered crystals of each mineral type in each parent rock ………………………………………………….... 134

Figure 29: Parent materials, post-reaction vessels, and dried products from closed system reactions ……………………………………………………………………………….… 137

Figure 30: A-C) Solid alteration products from the closed-system SMBxl run; D) Solid alteration products from the closed-system SMBgl run; E and F) Dried gel products from the SMBgl separated gel material; G) Solid alteration products from HN-BAS; H and I) Solid alteration products from CN-BAS; J) Solid alteration products from B-AND; K) Solid alteration products from AND; L) Dried gel products from AND; M) Solid products from RHY; N) Solid products from OBS; O) Dried gel products from OBS. Alg = alunogen; Hx hexahydrite; Pk = pickeringite; Fe-ox = Fe-oxide; An = anhydrite; Si = amorphous SiO2; Nat- j = natrojarosite; Fe-alun = Fe-bearing alunite; Nat-a = natroalunite; Fe/Al/Mg/mixed-sulf = unidentified sulfate phases ………………………………………………………………. 140

Figure 31: Major cation concentrations (ppb) in closed system fluids. Lack of data points for [K+] in SMBgl, HN-BAS, CN-BAS, and B-AND are due to concentrations below instrument detection limits. Due to their slow reaction rates, AND, RHY, and OBS parent materials have only experienced minor alteration, and the resulting geochemical fluids (shaded area) are therefore not considered to be indicative of equilibrium conditions ……………………. 141

Figure 32: Flow through SMBxl alteration products in SEM; Si – amorphous SiO2; An – anhydrite; Pk – pickeringite; Hx – hexahydrite; Nat-j – natrojarosite; Nat-alun - natroalunite ……. 146

Figure 33: Major element fluid chemistry evolution in flow through experiments; Lack of data points for [K+] in SMBxl are due to concentrations below instrument detection limits … 148

Figure 34: Flow through SMBgl Alteration products in SEM; Gl – unaltered glass; Si – amorphous SiO2; An – anhydrite; Pk – pickeringite; Hx – hexahydrite ……………………………... 150

Figure 35: Fluid composition evolution from flow through experiments. Total cation concentrations (sum of all cation concentrations in alteration fluid up to time X) are shown as Molar (M) concentrations, with totals increasing through time; Triangles = SMBgl; circles = SMBxl …………………………………………………………………………………. 157

Figure 36: Projected abundances of Fe-bearing secondary mineralogy for altered Fe-rich Martian basalts in closed systems (gray area). Blue triangles = our closed system altered terrestrial products; red stars = our closed-system altered SMB products; Orange circles = expected alteration products from unaltered Martian basalts (Lodders, 1998; H. Y. McSween et al., 2006; Ming et al., 2008; Rieder et al., 2004; Rubin et al., 2000) ……………………..… 161

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1. Introduction

Since the early days of Mars observation, humans have fantasized about life on the Red

Planet. Schiaparelli’s observation of “canali” on Mars (Manara and Trinchieri, 2011), and Percival

Lowell’s subsequent interpretation of those canali as signs of an advanced civilization (Manara and Wolter, 2011) were the initial sparks that fueled the flame in the search for life on Mars. Since the days of Schiaparelli and Lowell, our investigative capabilities have improved dramatically, but the burning question of whether life ever arose on Mars remains.

Despite its present-day arid state, more detailed orbital and in situ observation of Mars has resulted in extensive geomorphological and mineralogical evidence for a wetter past (e.g. Hynek et al., 2010; Mangold et al., 2012; Phillips et al., 2001; Siva and Sinha, 2017), though the “warm and wet” vs. “cold and icy” question is still hotly debated (Hynek, 2016). However, evidence for abundant crustal water early in Mars’ history, whether liquid or solid, has propelled the search for signs of microscopic life. Extremophile organisms thrive in terrestrial environments which we believe may be similar to those found on early Mars (Horneck, 2000). Of particular interest to many astrobiologists are hydrothermal systems and the thermophilic acid-tolerant species which flourish in them (e.g. Barns et al., 1994; Blank et al., 2002; Hugenholtz et al., 1998; Schulze-

Makuch et al., 2007), due to their potential to be some of the earliest forms of life on Earth (Martin et al., 2008), combined with the high preservation potential thanks to rapid precipitation of secondary minerals such as opaline silica (Allen et al., 2000; Water and Des Marais, 1993).

The investigation of relic Martian hydrothermal systems poses multiple challenges.

Secondary mineral assemblages which result from hydrothermal acid-sulfate alteration are similar to mineral assemblages which form in acidic saline lakes (Benison et al., 2007) and cold-based acid-sulfate alteration such as atmospheric acid fog (Banin et al., 1997). Distinguishing between

these formation mechanisms when selecting sites for detailed in situ rover investigations and confidently interpreting the geologic history upon arrival requires a deep understanding of the geochemical processes which may occur in these various environments, and the use of effective instrumentation methods for thorough mineral identification and site characterization.

This dissertation research addresses the following questions to improve our abilities in identifying relic hydrothermal systems on Mars and subsequently interpreting the geochemical conditions that existed in those potentially habitable systems:

1) What are the most effective in situ instrumentation methods to result in in-depth and

comprehensive observations of hydrothermal acid-sulfate alteration assemblages on Mars?

2) What influence does the iron content of the primary basalt have on secondary alteration

mineralogy?

3) Does alteration of iron-rich Mars basalt produce significantly different mineral

assemblages than those seen at common terrestrial analog sites?

The subsequent sections of this chapter will provide an overview of scientific observations of Mars and its geological history, including evidence for crustal water and hydrothermal systems. An overview of hydrothermal acid-sulfate alteration is also included, with details regarding relevant processes and products related to this work.

1.1. Observations of Mars: A History

1.1.1. Telescopic observations

The dawn of telescopic observation in 1609 brought a new wave of information to Mars science, which had until then been limited to naked-eye observations and calculations of planetary orbits. Initial observations of Mars were severely limited by optical imperfections in early telescope design. The first rough maps of Mars were drawn in 1636 by Francesco Fontana, though

they were little more than a disk shape. It wasn’t until 1659 that Christiaan Huygens resolved the rough shape of Syrtis Major through his significantly improved telescope. Additional details were then mapped by Cassini in 1666, and used to calculate the rotation of Mars about its axis. Further observation by Huygens led to the mapping of the north and south polar caps in 1672. More detailed observations of the polar caps and dark markings on Mars were done by Maraldi in the early 18th century, including noting changes in the size and shape of the polar caps through time.

Later observations by Herschel in 1783 calculated the axial tilt and center of the poles, and deduced the presence of a thin atmosphere. He also noted that Mars’ seasons were similar to Earth’s, just twice as long, and that variations in the bright polar regions were due to melting and freezing of material as the seasons changed.

By the end of the 1700s, Mars observation had turned primarily to distinguishing geographic features. Perhaps most notably, the observation team of Beer and Mädler conducted a systematic observation of Mars’ markings (including what was later named Sinus Meridiani, the present-day home of the MER Opportunity rover) in 1830, and determined them to be constant features, rather than atmospheric phenomenon as some previous observers had mused. Their observations resulted in the first Martian geographical map, which Mädler drew in 1840. Around the same time as Beer and Mädler’s observations, John Herschel surmised that the bright red areas on Mars were similar to Earth’s continents and the dark areas were likely seas. Over the following decades more detailed maps were drawn by Kaiser, and Lockyer, as well as the fateful 1877

Schiaparelli “canali” map, and Schiaparelli-inspired canal-filled map drawn by Percival Lowell in

1894.

1.1.2. Orbital and in situ spacecraft observations

The 1965 Mariner 4 flyby of Mars marked the dawn of a new era in Mars exploration. The

21 images that were returned by the spacecraft imaged ~ 1% of Mars’ surface in the vicinity of

Arcadia Planitia and Amazonis Planitia. These close-up images showed a heavily cratered planetary surface that was drastically different from the “canali,” continents, and seas of the telescopic era. Subsequent Mariner missions (Mariners 6 and 7) were sent to Mars in 1969 and imaged similarly cratered regions. It was not until the Mariner 9 mission reached Mars in 1971 that the true diversity of Mars’ surface was discovered. As the global dust storm which had engulfed the planet at the time of Mariner 9’s arrival settled, the spacecraft imaged the massive volcanoes of the Tharsis region poking through the haze. After the dust settled, the extensive canyon system which was later named Valles Marineris (after the Mariner missions themselves), the Martian south pole, features resembling river beds, and dunes were also photographed. This new wave of orbital imagery changed the perception of Mars from a heavily cratered and geologically dead planet to the geologically diverse and exciting planet we know it to be today.

The first in situ observations of Mars began in 1976 with the Viking 1 and 2 landers, as their counterparts simultaneously operated from orbit. The Viking 1 lander conducted observations of Chryse Planitia while the identical Viking 2 operated within Utopia Planitia. These landers acquired of thousands of images throughout their lifetimes, explored the Martian regolith and nearby rocks (Clark et al., 1982; Plumb et al., 1989), and conducted the still-hotly debated life- detection experiments which sought organic compounds in the soil (Klein, 1978).

A new phase of Mars observation began in 1997, with the arrival of the Mars Global

Surveyor (MGS) spacecraft as well as the Pathfinder lander and Sojourner rover. MGS carried a variety of instruments to gather remotely sensed data, including the Mars Orbiter Camera (MOC),

Mars Orbiter Laser Altimeter (MOLA), and the Thermal Emission Spectrometer (TES). The

combined data from MOC, MOLA, and TES provided high resolution imagery of the Mars surface

(MOC) (Malin et al., 2010), detailed topographical data (MOLA) (Smith et al., 2001), and a wealth of new information in the form of mineralogical composition and thermal inertia (TES)

(Christensen et al., 1998; 2001), which may serve as a proxy for grain size and help identify bedrock outcrops (Edwards et al., 2009; Ruff and Christensen, 2002). In situ exploration of a supposed fluvial plain in Ares Vallis was capable through Pathfinder’s visible imagery (Smith et al., 1997) and the Alpha Particle X-Ray Spectrometer (APXS) instrument (Foley, 2003), which was mounted on the Sojourner rover to measure elemental compositions of nearby rocks and soils.

The improved science payloads of the MGS spacecraft and Pathfinder lander (and accompanying Sojourner rover) provided significant insights into Mars’ geologic and hydrologic past. Increased image resolution of MOC allowed researchers to identify contemporary erosional gullies on crater and canyon walls, which were hypothesized to form through the seeping of liquid water (Malin and Edgett, 2000) or CO2 (Musselwhite et al., 2001) from the subsurface, and were observed to still be active (Malin et al., 2006). MGS-MOC images were also used in the identification of potential fluvial valley networks (Malin and Carr, 1999; Malin and Edgett, 2003), hinting at a warmer and wetter past. Pathfinder images of rocks in Areas Vallis also suggested a sedimentary source for some, while APXS data measured the composition of weathering rinds, which may also require water for their formation (Economou, 2001; Foley, 2003; Golombek and

Bridges, 2000; McSween et al., 1999; Morris et al., 2000; Wanke et al., 2001). Remotely-sensed mineralogical data from MGS-TES added further fuel to the argument for a wet early Mars, as hematite, carbonates, and sulfates – minerals or mineral groups which all require water for their formation – were all identified from orbit (Bandfield, 2002; Wray et al., 2016). Indeed, the

identification of hematite in Meridiani Planum was a major driver for future exploration of that area with the Opportunity rover (Golombek et al., 2003).

The early 2000s brought a renewed interest in Mars exploration thanks to earlier missions’ evidence for water. Several orbiters arrived at Mars in this time, including the Mars Odyssey

(2001), Mars Express (2003), and Mars Reconnaissance Orbiter (2006). In addition to orbital spacecraft, the hugely successful Mars Exploration Rovers (MER), Spirit and Opportunity, landed at Gusev crater and Meridiani Planum (respectively) in 2004. The MER investigations were followed by the next generation Mars Science Laboratory (MSL) Curiosity rover, which began investigating Gale crater in 2012. Details of relevant findings are described in sections 1.2 and 1.3.

1.1.3. Upcoming missions

Mars exploration is yet again on the cusp of a new era. The flood of data from the most recent generation of orbiters and rovers resulted in abundant mineralogical signs for aqueous activity in Mars’ past (Ehlmann and Edwards, 2014). This new influx of data led NASA to move on from the “follow the water” mantra of the early 2000s and adopt a new mission goal to seek signs of life (Beegle et al., 2007; Des Marais et al., 2008; Hamilton et al., 2015). Both NASA and

ESA-Roscosmos are preparing to send rovers with new-and-improved capabilities in the very near future. NASA’s Mars 2020 rover and the ESA-Roscosmos ExoMars rover are both scheduled to launch in 2020. Though landing site selection is still underway, the ExoMars team has narrowed the search to two possible locations: Mawrth Vallis and Oxia Planum (Bridges et al., 2016) while the Mars 2020 site selection has been narrowed to Jezero crater, northeast Syrtis Major, or returning to Columbia Hills in Gusev crater (Witze, 2017). These locations all contain high-value mineralogical targets such as phyllosilicates in Mawrth Vallis (McKeown et al., 2009), Oxia

Planum (Quantin et al., 2016), and Jezero crater (Goudge et al., 2015), phyllosilicates and sulfates

in Syrtis Major (Bramble et al., 2017), and SiO2 and sulfates in Gusev crater (Morris et al., 2008;

Ruff, 2015; Ruff et al., 2011; Squyres et al., 2008; Yen et al., 2008), which was previously explored by MER Spirit – all strong indicators of aqueous activity, and suggestive of potentially habitable environments at some point in Mars’ history.

1.2. The geology of Mars

1.2.1. A brief geologic history

Mars’ geologic history is divided into three major time periods: the Noachian (4.1 to 3.7

Ga), Hesperian (3.7 to 3.0 Ga), and Amazonian (3.0 Ga to present) (Figure 1). Little is known of the pre-Noachian time due to numerous large impacts and very little to no preservation of the rock record (Carr and Head, 2010). The dawn of the Noachian period is marked by the formation of the large Hellas impact basin at ~4.1 Ga (Frey, 2003) and continued high cratering rates (Hartmann and Neukum, 2001). In addition to impacts, the Noachian saw the formation of numerous valley networks (Hoke et al., 2011; Hynek et al., 2010), and deltas (Fassett and Head, 2005; Malin and

Edgett, 2003), the formation of phyllosilicates (and later sulfates) through aqueous processes

(Bibring et al., 2006; Murchie et al., 2009a), and extensive volcanic activity, including the formation of the large Tharsis region (Phillips et al., 2001; Robbins et al., 2011). This time in

Mars’ history has often been described as warm and wet (Craddock and Howard, 2002; Pollack et al., 1987; Squyres and Kasting, 1994) with evidence suggesting the presence of a large ocean filling the northern lowlands (Di Achille and Hynek, 2010). However, the “warm and wet” interpretation is still hotly debated (Hynek, 2016) with alternative models showing either intermittent wet phases (Halevy and Head, 2014; Kite et al., 2013), or cold and icy Southern

Highlands providing meltwater to form the fluvial and aqueous alteration products (Head, 2013).

The end of the Noachian (3.7 Ga) is marked by a significant decrease in cratering, aqueous

alteration, and fluvial processes (Carr and Head, 2010; Hartmann and Neukum, 2001) with volcanism continuing well into the Hesperian (Greeley and Guest, 1987; Scott and Tanaka, 1986;

Watters and Maxwell, 1986).

Figure 1: Timeline of the geologic activity on Mars from Carr and Head (2010)

Despite the decrease in valley formation rates, the Hesperian rock record still indicates the occurrence of water-based processes, including the formation of extensive outflow channels

(Leverington, 2011; Tanaka et al., 2005), intermittent valley networks (Fassett and Head, 2007;

2008, 2006, Gulick, 2001, 1998; Hynek et al., 2010), sulfate formation (Bibring et al., 2006;

Murchie et al., 2009a), and an increase in surface and subsurface ice (Hanna and Phillips, 2005;

Head and Pratt, 2001; Levy et al., 2009; Mangold et al., 2002). The large canyon system of Valles

Marineris also formed at this time, presumably as a result of stress fields related to formation of the Tharsis Rise (Blasius et al., 1977; Scott and Tanaka, 1986; Sharp, 1973). As time progressed into the Amazonian (3.0 Ga – present) the fluvial and volcanic rates decreased to a fraction of what is observed in the Noachian and Hesperian (Golombek et al., 2006), and geologic processes became dominated by the cryosphere and aeolian sediment transport we see today (Clifford, 1993;

Head and Marchant, 2003).

1.2.2. Volcanism and iron-rich basalts

The surface of Mars is composed primarily of basalt and basalt-derived sediments (Mustard et al., 2005), with volcanic constructs both large and small scattered across the planet (e.g. Wilson and Head, 1994). Despite the apparent present-day quiescence, the early history of Mars was shaped in large part by volcanism, with the formation of the vast Tharsis region (containing 5 of the largest volcanic structures in the solar system) nearly complete by the end of the Noachian

(Carr and Head, 2010; Kallenbach et al., 2001; Phillips et al., 2001), with episodic activity continuing up until ~100 – 10 Ma and possibly to the present (Hartmann and Berman, 2000;

Robbins et al., 2011). However, volcanism was not limited to only the Tharsis region. Volcanic structures such as paterae (collapse structures similar to ash-flow calderas), cinder cones, fissures

(Carr, 2006; Wilson and Head, 1994), and even potential “supervolcanoes” (Michalski and

Bleacher, 2013) are present across nearly all regions of Mars, and ash fall deposits have been suggested as an origin for the extensive finely layered mantling deposits in the area around Tharsis

(Hynek et al., 2003).

Although Mars appears to be a predominantly basalt-covered planet, Martian basalts are not compositionally identical to terrestrial basalts. Compositional analysis of relatively unaltered basaltic Mars meteorites (Shergottites LA 1 and QUE 94021) and in situ analysis of unaltered

basalts in Gusev crater and Meridiani Planum show a general enrichment of iron in Mars basalts relative to their terrestrial counterparts (Table 1).

Table 1: Iron abundances in unweathered Martian basalts and common terrestrial analogs

SiO2 FeOT Martian basalts Adirondack, Gusev 1 45.3 21.09 Backstay, Gusev 2 49.4 13.1 Irvine, Gusev 2 47.5 19.7 Algonquin, Gusev 2 41.9 20.9 Bounce Rock, Meridiani 3 50.8 15.6 Shergottite LA 1 4 49.1 21.2 Shergottite QUE 94021 5 47.9 18.5 Terrestrial basalts Hengill, Iceland 6 50.10 13.90 Krafla, Iceland 7 48.54 11.58 Kilauea, Hawaii – Pu’u O’o 8 52.32 10.77 Kilauea, Hawaii – 1974 flow 8 51.87 10.59 Cerro Negro, Nicaragua 9 49.73 9.70 Telica, Nicaragua 9 51.72 9.60 Momotombo, Nicaragua 9 54.40 9.10 Poás, Costa Rica 10 53.30 8.12 Turrialba, Costa Rica 10 54.14 7.59 Mauna Kea, Hawaii 11 49.74 6.02 1 back-calculated end member composition from McSween et al. (2006) 2 Ming et al. (2008) 3 Rieder et al. (2004) 4 Rubin et al. (2000); Hurowitz et al. (2005) 5 Lodders (1998) 6 Trønnes (1990) 7 Nicholson and Latin (1992) 8 Chemtob and Rossman (2014) 9 Hynek et al. (2013) 10 Black and Hynek (2017) 11 Morris et al. (2000)

This enrichment in Fe has been attributed to a combination of factors. Using only models of elemental condensation within the protoplanetary disk, Lewis (1972) concluded that Mars should have a Fe-rich mantle (FeO/(FeO+MgO) ≈ 0.5) and Fe-enriched crust relative to Earth. Later models and estimates using Shergottite compositions were in agreement with Lewis’ (1972) calculations, with an estimate of nearly double the Fe in Mars’ peridotite mantle compared to Earth

(Dreibus and Wänke, 1985). In addition to Fe-enrichment in the Mars mantle during planetary

accretion, crustal Fe is strongly influenced by the degree of partial melting during magma

production. Laboratory experiments with synthetic Gusev-composition basalts by Monders et al.

(2007) were in agreement with Dreibus and Wänke’s (1985) original mantle composition, and

concluded that partial melting on the order of 15 – 20% of the Fe-enriched mantle would result in

the observed Fe-rich Gusev basalt compositions.

1.2.3. Evidence for surface and subsurface water

1.2.3.1. Geomorphological evidence

Orbital and in situ investigations have provided a great deal of evidence for surface and

subsurface water on Mars through both geomorphological and mineralogical data. Orbital images

from MGS-MOC, and MRO’s High Resolution Imaging Science Experiment (HiRISE) and

Context (CTX) cameras as well as the Mars Odyssey Thermal Emission Imaging System

(THEMIS) and Mars Express High Resolution Stereo Camera (HRSC) have resulted in the

identification of extensive valley networks (e.g. Hynek et al., 2010), deltas (Di Achille and Hynek,

2010; Fassett and Head, 2005), and massive flood channels (Baker, 1978; Komar, 1979) (Figure

2).

Figure 2: Geomorphological evidence for flowing surface water; Left to right: valley networks, outflow channels, deltas, and gullies

1.2.3.1.1. Valley networks and deltas

Martian valley networks which resemble terrestrial dendritic fluvial systems were originally identified and mapped in Viking data by Carr (1995). However, this mapping was at low resolution due to the limitations of Viking imagery. The evolution of orbital data and the arrival of the MGS orbiter resulted in an improved global valley network map by Hynek et al.

(2010) showing far more numerous and dense valley networks across the mid-latitudes of Mars.

These valley networks were estimated to be from the Late Noachian – Early Hesperian (~3.8 – 3.6

Ga) boundary and are indicators of an extensive hydrologic system and sustained liquid water on the surface at that time (Hoke et al., 2011; Hynek et al., 2010).

In addition to valley networks, several deltas have been identified across the mid-latitudes of Mars (Di Achille and Hynek, 2010; Hauber et al., 2013; Hoke et al., 2014; Pondrelli et al.,

2008). These delta features have been likened to those on Earth, and interpreted as sites of fluvial deposition. Many deltas are located within craters, leading to interpretations of fluvial-fed crater lakes (e.g. Pondrelli et al., 2008).

1.2.3.1.2. Outflow channels

After the Mariner 9 mission captured images of expansive scour features on the surface of

Mars, it did not take long to recognize that these massive erosional networks are morphologically similar to the Channeled Scablands in eastern Washington state, and interpret them as the result of catastrophic flooding (Sharp and Malin, 1975). However, the driving process for the flooding remains a hotly contested debate – including evaporite dewatering (Montgomery and Gillespie,

2005), volcanically-driven melting of subsurface ice (Masursky et al., 1977), and overpressurization of aquifers by permafrost thickening (Carr, 1979). Regardless of the source of water, these features are indicative of large amounts of water flowing across the surface during the end of the Hesperian (Carr and Head, 2010).

1.2.3.1.3. Gullies

It was not until the MGS-MOC captured higher resolution images of crater and canyon

walls that fine scale gully features were observed at mid to high latitudes, primarily in the southern

hemisphere (Malin and Edgett, 2000). Malin and Edgett (2000) observed the gully features to be

devoid of craters, and depositing material on top of aeolian deposits – indicating very recent

activity. Indeed, repeated imaging of the gullies by MGS-MOC saw the formation of a new deposit,

suggesting these features are still active today (Malin et al., 2006). The association of gully heads

with specific stratigraphic units is suggestive of groundwater seepage as a fluid source and may be

an indicator of recent and present-day subsurface fluid flow (Malin and Edgett, 2000).

1.2.3.2. Mineralogical evidence

In addition to geomorphology, there is also extensive mineralogical evidence for water.

With the arrival of MRO’s Compact Reconnaissance Imaging Spectrometer for Mars (CRISM)

(Murchie et al., 2007) and the Mars Express Observatoire pour la Minéralogie, l'Eau, les Glaces et

l'Activité (OMEGA) (Bellucci et al., 2004) hyperspectral visible to near-infrared (VNIR)

reflectance spectrometers in the early 2000s, our ability to characterize Mars mineralogy greatly

improved. By measuring the reflected sunlight at various wavelengths (VNIR herein defined as

350 – 2500 nm unless otherwise noted), VNIR reflectance spectrometers may identify

characteristic absorptions that correspond with molecular vibrations within the surface material.

Absorptions within the VNIR range are particularly sensitive to H2O, OH, CO2, and Fe-bearing

minerals (Clark et al., 1990). This sensitivity makes CRISM and OMEGA exceptionally useful

tools in identifying and characterizing the global distribution of minerals such as phyllosilicates

(clays), sulfates, Fe-oxides/hydroxides, and hydrated SiO2 (Figure 3).

Figure 3: Global distribution of aqueous secondary mineralogy across Mars, from Ehlmann et al. (2014)

1.2.3.2.1. Phyllosilicates

Phyllosilicates have become a centerpiece of present-day Mars exploration (e.g. Bridges et al., 2015; Bristow et al., 2015; Poulet et al., 2014; Rampe et al., 2014), and form through aqueous alteration of basalts (e.g. Ehlmann et al., 2011a). Phyllosilicates have been identified across the surface of Mars within Noachian-age units, indicating extensive aqueous processes and environments early in Mars’ history (Bishop et al., 2008; Ehlmann and Edwards, 2014;

Greenberger et al., 2012; Le Deit et al., 2012; Loizeau et al., 2012; Michalski et al., 2010). These deposits appear as three different types – stratigraphically layered phyllosilicate units, phyllosilicates within crater rims and central peaks, and phyllosilicates within apparent depositional basins (Ehlmann and Edwards, 2014). Crater wall and central peak phyllosilicates likely formed through impact-induced hydrothermal alteration (Schwenzer and Kring, 2009) or are exposures of deep crust which were hydrothermally altered (Ehlmann et al., 2009), and

intracrater phyllosilicates through fluviolacustrine processes which ultimately resulted in their deposition within crater basins (Grindrod et al., 2012; Rampe et al., 2014). Perhaps more perplexing are the layered Fe/Mg- and Al-phyllosilicates, which have been identified in multiple locations, including Mawrth Vallis (Loizeau et al., 2007), Nili Fossae (Ehlmann et al., 2009),

Vallis Marineris (Murchie et al., 2009a), and throughout the southern highlands (Carter et al.,

2013). These layered deposits have been interpreted a number of ways, including alteration of dust or volcanic ash by melting ice/snow (Michalski et al., 2013), top-down leaching of basalt by surface water (Loizeau et al., 2012), and subsurface alteration by hydrothermal groundwater

(Ehlmann et al., 2011b).

1.2.3.2.2. Sulfates

Sulfates such as gypsum, alunite, and jarosite are common products of acidic alteration and have been identified at several locations on Mars (Ehlmann et al., 2016; Farrand et al., 2009;

Klingelhöfer et al., 2004; Weitz et al., 2013). Mixed sulfate-phyllosilicate intracrater deposits in

Terra Sirenum have been interpreted as evaporate deposits from acidic lakes (Glotch et al., 2010;

Swayze et al., 2008; Wray et al., 2009; Wray et al., 2008). Also located in Terra Sirenum is a ~10 km diameter deposit of alunite, kaolinite, and silica within Cross crater, which is suggestive of localized hydrothermal alteration (Ehlmann et al., 2016). Interior layered deposits within several sections of Valles Marineris contain both monohydrated and polyhydrated sulfates (sulfates with one or multiple H2O molecules in their structure) (Murchie et al., 2009a). The formation of these layered deposits has perplexed the scientific community, with proposed formation mechanisms ranging from fluviolacustrine to aeolian to volcanic (Nedell et al., 1987; Peterson, 1981), but is perhaps most likely a result of groundwater upwelling and cementation of aeolian sediment with evaporitic sulfates (Murchie et al., 2009). In addition to orbital sulfate detection, Martian sulfates

have also been identified in situ by the Spirit rover in Gusev crater (Wang et al., 2006b). The identification of Ca- and Mg-sulfates and silica in the soils near Home Plate were some of the first indications of prior hydrothermal activity in this region (Ming et al., 2008; Morris et al., 2008;

Schmidt et al., 2009, 2008; Squyres et al., 2007; Wang et al., 2006b). The sulfate deposits of Mars are diverse and widespread, and some of the strongest indicators for aqueous (and often acidic) processes occurring early in its geologic history.

1.2.3.2.3. Fe-oxides/hydroxides

Since the early days of Mars observation, the reddish hue of the surface has been a defining feature. We now know this to be due to a widespread global dust layer blanketing the surface in

Fe-oxides (Adams et al., 1986; Singer, 1982) and originates through acidic weathering of Martian basalts (e.g. Banin et al., 1997). Due to frequent mixing and redistribution through global dust storms, this dust layer is not indicative of the underlying geology (Kahn et al., 1990; Szwast et al.,

2006). However, there are many areas of Mars that are not blanketed in dust and still display characteristics of Fe-oxides/hydroxides within the bedrock itself (Christensen et al., 2001;

Ehlmann and Edwards, 2014; Lane et al., 2002; Murchie et al., 2009a; Poulet et al., 2007; Weitz et al., 2008).

Orbital detections of gray (coarse) hematite in Meridiani Planum by the MGS-TES instrument (Baldridge and Calvin, 2004; Lane et al., 2002) was a key factor in the selection of

Meridiani as the landing site for the Opportunity rover (Golombek et al., 2003). In situ investigation of Meridiani by Opportunity’s Mössbauer Spectrometer, which was designed specifically to detect iron-bearing minerals such as hematite (Klingelhöfer, 2004), confirmed the presence of hematite in the form of spherical concretions which were scattered throughout the

Meridiani bedrock and across the surface (Klingelhöfer et al., 2004). These hematite spherules

were interpreted to form as concretions within the jarosite-rich bedrock as acidic groundwater moved through the region (Squyres and Knoll, 2005).

While the Opportunity rover was investigating Meridiani Planum, its twin, the Spirit rover was also encountering Fe-oxides and hydroxides on the other side of the planet in Gusev crater

(Morris et al., 2008). Hematite and goethite were found in the Columbia Hills and Home Plate region of Gusev crater, along with Fe-sulfates and SiO2-enriched soils. Like those in Meridiani

Planum, these Fe-bearing deposits were indicative of strong acid alteration. However, in Gusev crater the combined mineralogy, geomorphology, and elemental enrichments point to a hydrothermal origin (Morris et al., 2008; Ruff, 2015; Ruff et al., 2011; Squyres et al., 2008; Yen et al., 2008).

1.2.3.2.4. Hydrated silica

SiO2-rich deposits have been detected both from orbit (Bandfield et al., 2013; Skok et al.,

2010) and in situ (Rice et al., 2010; Ruff et al., 2011; Squyres et al., 2008), and may serve as indicators for areas that have experienced extensive aqueous alteration and leaching due to the relative immobility of silica at low and neutral pH (e.g. Banin et al., 1997). In situ investigation by the Spirit rover encountered deposits of digitate opaline silica at Home Plate in Gusev crater which are strikingly similar in morphology to those from terrestrial hot spring environments (Ruff,

2015), and display similar characteristic absorptions within thermal emissivity data (Ruff et al.,

2011).

1.3. Hydrothermal systems on Mars

Many locations on Mars with mineral assemblages consisting of phyllosilicates, sulfates,

Fe-oxides, and hydrated silica have been identified using CRISM and OMEGA data (Carter et al.,

2013; Ehlmann et al., 2011b; Gendrin et al., 2005; S. L. Murchie et al., 2009a). Several formation

models have been proposed for these materials, including global-scale changes in the global-scale aqueous geochemistry (Bibring et al., 2006), acidic lakes (Baldridge et al., 2009), evaporation- driven playas (Andrews-Hanna et al., 2010), weathering of older phyllosilicate deposits (Altheide et al., 2010), and hydrothermal alteration (Morris et al., 2008; Ruff, 2015; Ruff et al., 2011;

Schmidt et al., 2009, 2008; Skok et al., 2010; Squyres et al., 2008; Thollot et al., 2012; Yen et al.,

2008). With our current understanding of hydrothermal alteration, Mars’ volcanic and aqueous history, and Martian mineralogy, there are some areas on Mars that show putative hydrothermal characteristics

1.3.1. Home Plate, Gusev crater

Although originally believed to be filled with lake sediments based on orbital observations

(e.g., Cabrol et al., 2003), in-situ observations at Gusev crater by Mars Exploration Rover (MER)

Spirit revealed basaltic plains (H. Y. McSween et al., 2006), pyroclastic deposits (Schmidt et al.,

2009, 2008; Squyres et al., 2007), and what is widely believed to be a preserved hydrothermal system at Home Plate (Morris et al., 2008; Ruff, 2015; Ruff et al., 2011; Squyres et al., 2008; Yen et al., 2008). Home Plate, a ~80 m diameter outcrop of altered volcanic tuff south of the Columbia

Hills, hosts SiO2 and sulfate-rich soils – indicative of high temperature, low pH alteration (2009,

2008). The association of these deposits with volcaniclastic materials suggests a hydrothermal origin. This interpretation is further supported by Alpha Particle X-ray Spectrometer (APXS) measurements showing the enrichment of sulfur, Si, and Ti, and depletion of highly mobile elements at the Fuzzy Smith float rock and in the Paso Robles class soils, a common pattern for basalts in contact with hydrothermal acid-sulfate waters (Ming et al., 2008; Yen et al., 2008). The high silica rocks of Home Plate have spurred heavy debate regarding the specifics of their origin; many have argued for a low fluid:rock hydrothermal origin (Hynek et al., 2013; Schmidt et al.,

2008; Yen et al., 2008), while Ruff et al. (2011) proposed a high fluid:rock hot spring system and detailed a terrestrial analog hot spring at El Tatio, Chile producing similar silica-rich deposits

(Ruff, 2015)

1.3.2. Eastern Coprates Chasma, Valles Marineris

Weitz et al. (2014) identified several light-toned outcrops in the walls of eastern Coprates

Chasma, where the Tharsis Rise intersects with Valles Marineris. CRISM data suggests these outcrops consist of Fe-rich amorphous hydrated silica, and their association with smectites, and the volcanically formed Tharsis Rise (Schultz and Tanaka, 1994) suggests a hydrothermal origin for these silica deposits

1.3.3. Noctis Labyrinthus, Valles Marineris

Several light-toned deposits are visible in multiple locations within Noctis Labyrinthus.

Analysis of CRISM data has revealed diverse mineralogies, including sulfates, hydrated silica,

Fe/Mg smectites, and Al-rich phyllosilicates. Multiple formation mechanisms are possible, including acidic aqueous activity and alteration of pre-existing material, and acid sulfate hydrothermal alteration of basalt and volcanic ash (Thollot et al., 2012, 2011; Weitz et al., 2011).

Mineralogical variations between locations within Noctis Labyrinthus imply that formation mechanisms were likely localized, and did not act on Noctis Labyrinthus as a whole (Weitz et al.,

2010), which lends support for the hydrothermal alteration formation mechanism, as these processes a generally more localized than aqueous processes.

1.3.4. Cross crater, Terra Sirenum

A large (10 km × 5 km) alunite-bearing deposit was identified in Cross crater (Ehlmann et al., 2016) using spectral data acquired with CRISM (S. L. Murchie et al., 2009b) – one of the few, and the largest, alunite deposits identified on Mars thus far. CRISM spectra gathered within Cross

crater show large deposits of alunite, kaolinite, and montmorillonite/opaline silica. This mineralogy is markedly different from the surrounding areas, and indicative of localized aqueous acid-sulfate alteration within the crater limits. Ehlmann et al. (2016) offer multiple possible formation mechanisms but note that upwelling of magma-heated acid-sulfate waters is the best match for the observed deposits.

1.3.5. Nili Patera, Syrtis Major

Skok et al. (2010) described several hydrated silica deposits around the summit area of a cone within the Nili Patera caldera. A broad fan shaped deposit is identifiable on the west flank, as well as several smaller circular and semicircular deposits on the patera floor, and thin layers visible in outcrop on the west side of the cone. The association with the seemingly volcanic cone strongly suggests a hydrothermal formation for these deposits.

1.4. Acid-sulfate alteration: Processes and products

Hydrothermal acid-sulfate alteration is a result of the interaction between aqueous H+ and primary minerals, wherein the H+ attacks and breaks down the original silicate mineral. This process releases cations to the hydrothermal fluid, and produces new secondary mineral phases.

H+ is abundant and readily available in hydrothermal environments due to the oxidation of magmatic sulfur dioxide (SO2) into sulfuric acid (H2SO4) and reduction to hydrogen sulfide (H2S)

(Equation 1), as well as further oxidation of the H2S into H2SO4 (Equation 2). The resulting H2SO4

3 -1.9 + will readily deprotonate (Ka1 = 10 ; Ka2 = 10 ) and contribute additional H to the system

(Equations 3 and 4).

4푆푂2 + 4퐻2푂 → 3퐻2푆푂4 + 퐻2푆 (1)

퐻2푆 + 2푂2 → 퐻2푆푂4 (2)

+ − 퐻2푆푂4 → 퐻 + 퐻푆푂4 (3)

− + 2− 퐻푆푂4 → 퐻 + 푆푂4 (4)

Hydrothermal acid-sulfate alteration is characterized by strong acid leaching of primary minerals, and the formation of a high silica enrichment zone in the areas proximal to fumarolic vents and hyperacidic crater lakes. An alunite-dominated alteration zone often surrounds the silica-enriched core, with a kaolinite-dominated zone in the distal regions (Heald et al., 1987;

Pirajno, 2010). Also common are Mg-, Fe-, and Ca-sulfates such as gypsum, jarosite, natrojarosite,

Fe-rich natroalunite, hexahydrite, and kieserite, Fe- and Mg- smectites such as nontronite and montmorillonite, elemental sulfur, and Fe-oxides (Table 2) (Bishop et al., 2004; Ehlmann et al.,

2012; Hynek et al., 2013; Marcucci et al., 2013; McCollom et al., 2013a; Rodríguez and van

Bergen, 2015).

Table 2: Common hydrothermal alteration products Mineral Composition Hydrated silica SiO2 * n H2O Sulfates Gypsum CaSO4 * 2 H2O Alunite/Natroalunite (K,Na)Al3(SO4)2(OH)6 Jarosite/Natrojarosite (K,Na)Fe3(SO4)2(OH)6 Hexahydrite MgSO4 * 6 H2O Phyllosilicates Nontronite Na0.3Fe2((Si,Al)4O10)(OH)2 * 4 H2O Montmorillonite (Na,Ca)0.33(Al,Mg)2(Si4O10) Kaolinite Al2Si2O5(OH)4 Fe-Oxides Hematite Fe2O3 Goethite FeO(OH)

Hynek et al. (2014, 2013) conducted a survey of several sites across Nicaragua, Costa Rica,

Iceland, and Hawaii to examine the relationship between pH, temperature, water/rock ratio, parent lithology, and alteration mineralogy. Results suggest elemental enrichment in the parent lithology provides an abundance of cations for sulfate formation (e.g., alteration of Fe-rich parent rock would result in the formation of Fe-sulfates), and has significant influence over the resulting secondary mineralogies. Additionally, Hynek et al. (2013) and Marcucci et al. (2013) observed

phyllosilicates forming in low pH environments with high fluid/rock ratios, although these materials are commonly interpreted as indicators of neutral to alkaline environments (Bibring et al., 2006; Chevrier et al., 2007).

1.5. Dissertation research

Proper characterization of mineralogy is an essential part of geologic interpretation.

Hydrothermal alteration products often exist in intimate mixtures, and vary widely across a variety of spatial scales due to differing pH, temperature, and fluid/gas chemistries. These characteristics require that we develop a detailed understanding regarding the environmental controls on resultant mineral mixtures, and how these various combinations appear in different instrument data sets.

Mars rovers carry with them a wide range of analytical tools to aid in the interpretation of Mars’ geologic history. However, many instruments are best suited to detect specific materials, and miss others. Often, multiple instruments must be utilized to develop a comprehensive understanding of a sample. Additionally, our interpretations of Mars data, and the resulting inferred geologic histories, are based on the results of terrestrial analog studies and laboratory simulations. Very few of which, thus far, take into account large differences in basalt composition between Earth and

Mars. This is in spite of the fact that previous work (Hynek et al., 2014, 2013; Marcucci et al.,

2013; Schwenzer and Kring, 2013; Tosca et al., 2004) suggests parent lithology strongly influences the mineralogical composition of alteration products.

In this dissertation, I analyze a wealth of samples acquired at terrestrial sites using multiple lab- and field-based instruments. Instruments utilized in this work (VNIR, XRD, and Raman) have analogous Mars rover counterparts. I use these datasets, along with environmental conditions measured in the field, to develop a detailed picture of each of my Icelandic field sites and the hydrothermal alteration products that are present across varied environmental conditions.

Mineralogical data from Costa Rica, Iceland, and Nicaragua is also used to determine what alteration products and mixtures may be identified using different analyses, as well as identifying instrumentation blind spots, and determining which method, or combination of methods results in the most thorough and accurate identification.

Natural systems have many uncontrolled variables; therefore, I conduct a suite of laboratory experiments to aid in interpretation of the field data and application to Mars’ relic hydrothermal systems. For my experimental work, I simulate both flow-through and closed hydrothermal systems with 1M H2SO4 acid and simulated Mars basalt, as well as a suite of terrestrial volcanics which range in composition from high-Fe basalt to rhyolite to systematically assess the effects of parent lithology on secondary mineralogy. Unlike previous works, this study simulates hydrothermal alteration in both crystalline and glassy synthetic Mars basalt based primarily on in-situ compositional data from the Spirit rover. Previous studies (Baker et al., 2000;

Horgan et al., 2017; Hurowitz et al., 2005; Peretyazhko et al., 2018; Tosca et al., 2004; Yant et al.,

2016) have examined acid sulfate alteration in synthetic Mars basalt or high-Fe terrestrial basalts, however, none of these experiments combined extended duration reactions, high-Fe Mars composition basalts of both a glassy and crystalline nature, and hydrothermal conditions, as I have done here.

If we want to search for signs of ancient life on Mars, one likely target is relict hydrothermal systems. Terrestrial hydrothermal systems support communities of microbial life (e.g. Blank et al.,

2002; Hugenholtz et al., 1998), and the rapid precipitation of minerals such as hydrated silica results in preservation of biosignatures in the form of microbial morphologies (e.g. Allen et al.,

2000; Water and Des Marais, 1993). In our search for Martian environments which may have been habitable, and may even contain preserved signs of life, we must use the remaining mineralogical

deposits to interpret the geochemical conditions at the time that environment existed. Factors such as pH, temperature, and fluid abundance have a significant impact on whether a site may support microbial life (e.g. Preston and Dartnell, 2014), and the minerals which remain in the rock record are the tools that allow us to interpret those conditions millions of years later. However, the mineralogy of hydrothermal systems is highly variable on the fine scale, and must be assessed in- situ to have any confidence in our interpretations. Therefore, it is vital that we improve our understanding of alteration processes, as well as our interpretative capabilities using in-situ rover data. All previous interpretations of hydrothermal alteration on Mars have been made using either terrestrial field analogs, or experimental setups, with have either not matched the basalt compositions we observe on Mars, did not simulate hydrothermal conditions, or only used short- duration experiments. This study continues to explore the mineralogical composition of terrestrial hydrothermal sites, as well as the effect of Mars basalt composition on alteration products.

2. Characterization of terrestrial hydrothermal alteration products with Mars analog instrumentation: Implications for current and future rover investigations

Note: This chapter was published in the journal Icarus in October 2017 as: Black, S. R., and Hynek, B. M. (2017), “Characterization of terrestrial hydrothermal alteration products with Mars analog instrumentation: Implications for current and future rover investigation.” Icarus. It is cited in the reference list and other chapters as “Black and Hynek, 2017.” Sections and figures have been re-numbered to match the formatting of the dissertation. Acknowledgements have been combined at the beginning of the dissertation. All work was completed by SRB.

Abstract: Interpretation of Martian geology relies heavily on our understanding of terrestrial analog deposits and our ability to obtain comprehensive and accurate mineralogical compositions. Many previous studies of terrestrial hydrothermal deposits relied on limited datasets and/or did not use instruments analogous to those deployed on Mars. We analyzed 100 hydrothermally altered basalts from Costa Rica, Nicaragua, and Iceland with Mars analog Visible to Short Wave Infrared (VSWIR) spectroscopy, X-Ray Diffraction (XRD), and Raman laser spectrometry. Alteration mineralogy consisted of amorphous and crystalline SiO2 (cristobalite, tridymite, quartz), Ca/Al/Fe/Mg-sulfates (gypsum, anhydrite, alunite, jarosite, hexahydrite, alunogen), Fe-, Ti-, and Mg-oxides/hydroxides (hematite, goethite, anatase/brookite, brucite), elemental sulfur, and phyllosilicates (montmorillonite, kaolinite). Results indicate VSWIR is best suited for identification of X-ray amorphous materials such as hydrated SiO2 and phyllosilicates, while XRD is best utilized for highly ordered crystalline materials such as sulfates, crystalline SiO2 polymorphs, elemental sulfur, and Mg-hydroxides identification. Surprisingly, XRD had the lowest identification rates for Fe-oxides/hydroxides (42% compared to 61% and 75% for VNIR and Raman, respectively), and nearly equal identification rates as VSWIR for kaolinite (76% for VSWIR, 71% for XRD). Identification of phyllosilicates in XRD, while possible, is not as effective as VSWIR without extensive sample preparation. Our observed identification rates may be attributed to the relative abundance of materials – Fe-oxides/hydroxides being present as surface coatings, the presence of large amounts of kaolinite in some samples, and an increased particle size for kaolinite relative to other clays. Elemental sulfur and Fe- and Ti-oxides/hydroxides were more readily identified with Raman. With NASA’s current focus on habitability, hydrothermally altered areas – which we know to host a wide range of microbial life here on Earth – are of high interest and it is likely that future rovers will encounter similar mineral assemblages. Therefore, future rovers would benefit from using a combination of these methods and expanding the VSWIR sampling range to the full 300 – 2500 nm to conduct a comprehensive mineralogical investigation.

2.1. Introduction

2.1.1. Hydrothermal alteration on Mars

Volcanism has played an active role throughout much of Mars’ geologic history – the remnants of which may be seen today in the form of volcanic edifices both large and small (e.g.

Crumpler et al., 1996; Wilson and Head, 1994), lava flows (e.g. Greeley and Crown, 1990;

Mouginis-Mark and Rowland, 2008), pyroclastic and airfall deposits (e.g. Brož and Hauber, 2012;

Gregg and Farley, 2006; Hynek et al., 2003; Squyres et al., 2007), and ancient hydrothermal systems (e.g. Skok et al., 2010; Squyres et al., 2008). Stratigraphic relationships and the crater record suggests volcanism was one of the dominant geologic processes early in Mars’ history (>3.5

Ga), and continued – localized to the Tharsis region – until 100-200 Ma (Neukum et al., 2004;

Robbins et al., 2011; Werner, 2009), and potentially up to < 10 Ma (Hartmann and Neukum, 2001).

Furthermore, impact crater counts and isotopic age dating of SNC meteorites indicate that Mars’ volcanism continued sporadically up until at least 100 Ma (Nyquist et al., 2001), and several studies have shown the potential for long-lived impact-induced hydrothermal systems on Mars

(Marzo et al., 2010; Schwenzer and Kring, 2013, 2009) in addition to volcanic systems. Observed

Martian mineralogy (hydrated sulfates, SiO2, and phyllosilicates) (Bishop, 2005; Ehlmann et al.,

2011a; Milliken et al., 2008; Mustard et al., 2008) and geomorphological features such as valley networks (Hynek et al., 2010; Wilson et al., 2016), deltas and paleolake deposits (Di Achille and

Hynek, 2010; Hynek et al., 2015; Pondrelli et al., 2008), and outflow channels (Andrews-Hanna and Phillips, 2007; Wilson et al., 2004) require large amounts of crustal water to have been present throughout much of Mars’ history. This combination of heat from volcanism and impact cratering and the presence of crustal water should have created hydrothermal systems on Mars through time, and we should expect to see evidence of these ancient hydrothermal systems preserved on Mars today.

Hydrothermal systems are of great interest for the astrobiology community. Investigations of numerous terrestrial systems show the ability for extremophile life to thrive in these environments (e.g. Blank et al., 2002; Walker et al., 2005), and may be similar to those that

harbored the earliest forms of life (Reysenbach and Cady, 2001). In addition to hosting robust microbial communities, rapid mineral precipitation – particularly that of SiO2 in the form of opal or siliceous sinter – in hydrothermal systems provides an excellent medium for preserving biosignatures (e.g. Al-Hanbali et al., 2001; Walker et al., 2005).

In the search for signs of ancient life on Mars, one likely target is relict hydrothermal systems. Hydrothermal systems have been implicated in the origin of life on Earth (Reysenbach and Cady, 2001) and can provide the necessary materials and energy sources for life (Martin et al.,

2008). To identify potentially habitable locations on Mars, we seek the mineralogical signatures that are associated with these hydrothermal systems. Many locations on Mars with mineral assemblages consisting of phyllosilicates, sulfates, and hydrated silica have been identified using the Mars Reconnaissance Orbiter’s (MRO) Compact Reconnaissance Imaging Spectrometer for

Mars (CRISM) and the Mars Express’s (MEx) Observatoire pour la Minéralogie, l’Eau, les Glaces et l’Activité (OMEGA) data (Carter et al., 2013; Ehlmann et al., 2011a; Gendrin et al., 2005;

Murchie et al., 2009a). Deposits include hydrated silica around the summit area of a cone within the Nili Patera caldera (Skok et al., 2010), light-toned deposits of sulfates, hydrated silica, Fe/Mg- smectites, and Al-rich phyllosilicates within Noctis Labyrinthus (Thollot et al., 2012; Weitz et al.,

2011), and light-toned outcrops of Fe-rich amorphous hydrated silica, smectites in the walls of eastern Coprates Chasma (Weitz et al., 2014), and large alunite deposits in Cross Crater (Ehlmann et al., 2016).

In addition to orbital data, observations by the Sprit rover found soils near Home Plate, a subcircular unit resembling a volcanic tuff approximately 80 m in diameter (Schmidt et al., 2009,

2008; Squyres et al., 2007), containing hydrated ferric iron sulfates, and amorphous silica, as well as digitate opaline silica nodules in the Elizabeth Mahon outcrop. These materials are consistent

with the interpretation of low pH hydrothermal alteration in this area (Morris et al., 2008; Ruff,

2015; Ruff et al., 2011; Squyres et al., 2008; Yen et al., 2008).

Recent detections of tridymite by the Curiosity rover in Gale Crater suggest a history of hydrothermal activity either within Gale Crater or in the surrounding catchment area (Morris et al., 2016). It is therefore plausible that hydrothermal deposits from the surrounding region are represented in the Gale Crater rock record due to subsequent erosion and deposition.

2.1.2. Instrumentation

Remote investigations of planetary bodies often use multiple instrumentation methods in tandem to explore a target such as the relict hydrothermal system at Home Plate in Gusev crater

(Morris et al., 2008; Ruff, 2015; Ruff et al., 2011; Squyres et al., 2008; Yen et al., 2008). This study utilizes three investigative methods that are deployed on past, current, and/or future Mars rovers to investigate terrestrial hydrothermal deposits. Although multiple instruments are often used to explore a target, it is important to consider the strengths and weaknesses of each instrument being used to ensure that no component is missed in analysis. Our work explores the effectiveness of Visible to Short Wave Infrared (VSWIR) reflectance spectroscopy, X-ray Diffraction (XRD), and Raman laser spectroscopy when working with common hydrothermal acid-sulfate secondary mineralogies. Our aim is to identify what minerals are often not identified by each method, so that future investigations may tailor their science payloads for maximum scientific return.

2.1.2.1. Visible-Near Infrared (VNIR) and Short Wave Infrared (SWIR) Reflectance Spectroscopy

VNIR/SWIR (herein referred to as “VSWIR”) reflectance spectroscopy utilizes reflected light from 350 to 2500 nm to identify the mineralogical composition of a material. Each mineral has a characteristic pattern of absorptions, which is determined by its chemical composition and bonding within the crystal structure. VSWIR is sensitive to the presence of transition metals,

OH/H2O, CO3, and PO4 molecules. The most common absorption bands which aid in the identification of hydrothermal alteration minerals are those which correspond to electronic transitions of Fe2+ and Fe3+, sulfur conduction bands, combinations and overtones of the fundamental H2O and OH bend and stretch vibrations, as well as combinations and overtones from fundamental metal-OH bend and stretch vibrations, such as Al-OH, Mg-OH, and Fe-OH. A thorough overview of vibrational and electronic spectroscopy is given in (Hunt, 1977) and (Clark et al., 1990). VSWIR can be used at any scale – from km to m – and can be deployed orbitally and used as a remote sensing tool, or on the ground to observe small scale features. VSWIR spectrometers can utilize both passive (solar) and active (e.g. an internal lamp) illumination, which makes these versatile instruments for a variety of applications. Although VSWIR is a useful tool in mineralogical characterization, by its nature reflectance spectroscopy is only sampling the top few microns of a sample (Buckingham and Sommer, 1983), and may not be representative of the bulk composition. Additionally, VSWIR is primarily used as a qualitative mineralogical characterization tool and requires significant processing and modeling with radiative transfer methods to obtain quantitative results (Stack and Milliken, 2015).

VNIR broad band sensors are regularly utilized on Mars rovers, as they return data rapidly, do not require large amounts of additional processing, and are not limited to contact analysis and can therefore be done in situ or remotely. Remotely sampling a region with VSWIR results in large scale mineralogical mapping of a region, and can give a wealth of compositional data in one scene, as opposed to the detailed but spatially limited point-analysis that is done by contact instruments such as X-ray diffraction (XRD) spectrometers. JPL defines VNIR as the 300 – 1000 nm spectral range, however, most laboratory-based instruments cover the full 350 to 2500 nm

VSWIR range. For consistency, we will use the term VSWIR to mean the 350 to 2500 nm range

and VNIR as defined by JPL (300 – 1000 nm). In-situ VNIR Mars data has been gathered using

Mars Exploration Rover (MER) PanCam, as well as The Mars Science Laboratory (MSL)

MastCam and ChemCam’s passive Laser-Induced Breakdown Spectroscopy (LIBS) function

(Table 3). The MER PanCam measures reflected light from 400 to 1100 nm on 11 channels (Bell et al., 2003). Like the MER PanCam, MSL’s MastCam also measures 11 channels, but from 400 to 1000 nm (Bell et al., 2012; Malin et al., 2010). The MSL ChemCam passive LIBS measures from 400 to 840 nm, with a <1 nm/band resolution over 2048 channels (Johnson et al., 2015). All instruments have spectral windows that are best suited for the identification of Fe- oxides/hydroxides – a common material on the Martian surface. It is important to note, however, that dust is ubiquitous on the surface of Mars and will obscure the spectral signature of the rocks over which it lies.

Upcoming rovers also have VNIR/VSWIR instruments as part of their payload (Table 3).

The Mars 2020 rover will have a spectrometer as part of the SuperCam and Mastcam-Z instruments

(Clegg et al., 2014; Maurice et al., 2015; Wiens et al., 2016). The ExoMars rover will have the

PanCam (Harris et al., 2015), MicroOmega Infrared Spectrometer (Leroi et al., 2009), the Infrared

Spectrometer for ExoMars (ISEM) (Korablev et al., 2014), and the Mars Multispectral Imager for

Subsurface Studies (Ma_MISS) (Coradini et al., 2001), which will be mounted on the side of a 2 meter long drill probe, and capable of taking spectral measurements of the subsurface from within boreholes (Bost et al., 2015).

Table 3: Instrument comparison. Instruments in bold are terrestrial functional-equivalents used in this study. Plain text represents instruments currently deployed on Mars. Italic text represents instruments planned for upcoming Mars missions. VNIR/VSWIR Instrument Spectral window (nm) Spectral resolution Spatial resolution Area sampled/FOV TerraSpec 4 and HALO 1 350 – 2500 6 nm/band N/A – one spectra for FOV 1 cm diameter MER: PanCam 2 400 – 1000 11 spectral bands 2.8 cm/px at 100 m 1 FOV MSL: MastCam 3 400 – 1000 11 spectral bands Up to 450 µm/px at 2 m 6.8 x 5.1 FOV MSL: passive ChemCam 4 400 – 840 0.61 nm/band N/A – one spectra for FOV ~1.0 mm diameter Mars 2020: MastCam-Z 5 400 – 1000 11 spectral bands Up to ~0.15 mm/px at 2 m 17 FOV Mars 2020: SuperCam 6 400 – 900, 1300 – 2600 > 248 bands N/A – one spectra for FOV 1.3 mm at 2 m ExoMars: PanCam 7 440 – 1000 11 spectral bands ~1cm/px (1024 x 1024) 38.3 FOV ExoMars: Ma_MISS 8 400 – 2200 20 nm/band 100 µm/px 0.1 mm diameter ExoMars: MicrOmega 9 500 – 3500 < I nm/band (900-3500 nm) 20 µm/px, 5 mm FOV 5 mm diameter ExoMars: ISEM 10 1150 – 3300 < 9 nm/band N/A – one spectra for FOV 1 FOV XRD Instrument Diffraction window (2) Resolution (2) Ka source Terra XRD 11 5 – 55° 0.25 ° Cu-Ka MSL: CheMin 12 5 – 50° < 0.35° Co-Ka ExoMars: Mars-XRD 13 6 – 65° not stated 55Fe- Ka Raman Instrument Laser type Spectral window (cm-1) Spot size (µm) High Res. Mapping Horiba LabRam HR Evolution 14 532 nm (green) 100 – 1700 cm-1 5 µm (10x), 2 µm (50x), 1 µm (100x) Yes Mars 2020: SHERLOC 15 248 nm (DUV) 810 – 3500 cm-1 50 µm Yes Mars 2020: SuperCam 6 523 (green) 0 – 4000 cm-1 1.3 mm at 2 m “Scan mode” ExoMars: RLS 16 532 nm (green) ~150 – 3800 cm-1 50 µm Yes

1: ASD Inc. (2015a, 2015b) 2: Bell et al. (2003) 3: Malin et al. (2010), Bell et al. (2012) 4: Johnson et al. (2015) 5: Clegg et al. (2014) 6: Clegg et al. (2014), Maurice et al. (2015), Wiens et al. (2016) 7: Griffiths et al. (2006); Cousins et al. (2012) 8: Coradini et al. (2001), Bost et al. (2015) 9: Leroi et al. (2009) 10: Korablev et al. (2014) 11: Olympus Scientific Solutions (2014) 12: Sarrazin et al. (2005), Blake et al. (2012) 13: Hill et al. (2011), Marinangeli et al. (2011) 14: Ellison (2017)

15: Clegg et al. (2014), Beegle et al. (2015)

2.1.2.2. X-Ray Diffraction (XRD)

XRD may be used to determine the bulk quantitative mineralogy of a sample. In this technique, powdered samples are bombarded by x-rays and diffracted, which results in peaks in the diffraction pattern. Peak location is determined by the crystalline structure of the minerals composing the sample. XRD analysis is generally considered to be the most complete compositional assessment (Ehlmann et al., 2012), if only one contact analysis can be done. When used in conjunction with additional data sets, XRD provides a valuable baseline against which other compositional analysis may be compared such as ground truthing remotely-sensed VSWIR data (e.g. Swayze et al., 2014).

While XRD is often used for bulk compositional analysis for materials present at or above the ~3 wt % detection limit (Blake et al., 2012), it is best for identifying coarser, highly-crystalline materials. If a minor phase, phyllosilicate identification can require several additional processing steps such as sieving and centrifugation to separate out the fine particulates, oriented mounts to amplify the 001 basal reflection that is preferred in phyllosilicates, and ethylene glycol washes to observe the swelling – or lack of – to distinguish between various types of phyllosilicates

(Beckerman, 2016; Środoń, 2006). Diffraction patterns of untreated samples will be dominated by peaks from the larger, highly-crystalline particles due to their increased diffraction relative to smaller grained materials (McCusker et al., 1999). Crystalline minerals which are present in trace quantities below the ~3 wt % detection limit are unlikely to appear in the diffraction pattern, and must be identified through a different method, such as VSWIR. Poorly crystalline or nanophase materials such as nanophase Fe-oxides are XRD amorphous, and often to not appear in diffraction patterns even when present above detection limits (Swayze et al., 2000).

XRD is currently utilized on Mars with the CheMin X-ray Diffraction Instrument on the

MSL Curiosity Rover (Table 3) (Blake et al., 2012; Sarrazin et al., 2005). CheMin’s XRD utilizes

Co-K x-rays, which offers improved peak resolution, and identification of peaks at low 2 values

– often those which correspond to phyllosilicates – due to the overall shift to larger 2 values and larger spacing between peaks (Harris and White, 2007; Stahly, 2012). CheMin measures from 5 to

50° 2, with a resolution of <0.35° 2. This sampling range covers a wide range of peaks, and is capable of identifying all known materials that may be present on the Martian surface if they are present in concentrations above XRD detection thresholds of ~3 wt % (Blake et al., 2012; Swayze et al., 2014). Future XRD investigation of Mars includes the European Space Agency’s ExoMars rover, which has the Mars-XRD instrument planned for the science payload and is scheduled to launch in 2020 (Table 3) (Hill et al., 2011).

2.1.2.3. Raman Laser Spectroscopy

Raman spectroscopy observes the shift in wavelength when a laser of a known wavelength is directed at a sample. The incident ray is scattered, primarily through elastic Rayleigh scattering.

However, a small number of photons will experience inelastic Raman scattering, and are absorbed and subsequently re-emitted at a different wavelength than the original laser. The shift in wavelength is determined by molecular vibrations similar to those that create VSWIR absorptions.

In the case of Raman, however, energy is emitted as peaks indicating various fundamental vibrations, rather than absorbed as in VSWIR. These peaks are characteristic of the elemental composition and molecular structures present in the sample, and allow for determination of mineralogical composition. Raman spectroscopy is particularly sensitive to OH, H2O, CO2, SO4, and Si-O bonds (Vandenabeele, 2013), allowing for the identification of many common hydrothermal minerals.

Raman spectrometers must provide their own active illumination source in the form of a

laser. Common laser wavelengths that are used for Raman are 532 nm (green) and 248 nm (deep

UV). 532 nm is the most common wavelength for mineralogical investigations, and is limited primarily by the large amount of fluorescence that is created when the laser contacts amorphous or poorly crystalline materials such as hydrated SiO2 or phyllosilicates. Deep UV Raman is most often used when investigating organics, and may also be used to identify phyllosilicates such as kaolinite, as these materials do not fluoresce when in contact with the deep UV laser. Although

Raman laser spectrometers can sample materials up to several meters away, there is a practical distance limit for each instrument which is determined by the laser and detector specifications.

Several Mars analog studies have been conducted using Raman spectrometers. Wang et al. (2006a, 1999) and Israel et al. (1997) demonstrated the utility of Raman for mineralogical characterization of Mars-like basalts and various types of alteration rinds, as well as the hydration states of Mg-sulfates. Successful mineralogical analyses were performed on varnished basalts from the Lunar Crater Volcanic Field (Israel et al., 1997), hydrothermally altered basalts from the

Keweenawan North Shore Volcanics in Minnesota (Wang et al., 1999), and laboratory-produced

Mg-sulfates (Wang et al., 2006a) – all of which are mineralogies likely to be encountered on Mars.

There are no Raman spectrometers currently operating on Mars; however, the upcoming

Mars 2020 and ExoMars rovers both have Raman spectrometers as part of the instrument payload

(Table 3). The Mars 2020 rover will have both the SuperCam – carrying a 532 nm Raman laser which will be capable of sampling outcrops up to 12 m away (Clegg et al., 2014) – and the

Scanning Habitable Environments with Raman & Luminescence for Organics & Chemicals

(SHERLOC) – a deep UV Raman laser, primarily suited for the detection of organics (Beegle et al., 2015; Clegg et al., 2014) – as part of its payload. The ExoMars rover will have the Raman

Laser Spectrometer (RLS) – another 532 nm Raman laser, similar to the SuperCam Raman on

Mars 2020 (Lopez-Reyes et al., 2013; Rull et al., 2011).

2.1.3. Characterization of terrestrial hydrothermal deposits

Hydrothermal systems are highly variable on the fine scale, and may produce similar mineral assemblages as those found in acidic lakes (Baldridge et al., 2009). While remotely-sensed data may be useful for big picture questions and aid in distinguishing between these depositional environments, in-situ observations provide small-scale details and additional insight into the geological processes that have occurred in a region – particularly those that have been subsequently weathered and eroded and/or are no longer easily observable from orbit. By combining regional scale orbital observations such as CRISM mineralogy with fine scale rover observations we can develop a deeper understanding of the geologic history of a location. Mars rovers carry with them a wide range of analytical tools to aid in geologic interpretation, but many instruments are best suited to detect specific materials, while missing others. Often, multiple instruments must be utilized to develop a comprehensive understanding of a sample. Thorough characterization of a sample is necessary for interpreting the geochemical conditions of formation, as various minerals such as phyllosilicates and sulfates form in specific ranges of pH and temperature. Failure to identify a mineral that is present in a sample may lead to missing critical indicators of geochemical conditions and result in incorrect interpretations.

A number of studies have investigated hydrothermal alteration products – from those formed from impact-induced hydrothermal alteration (Hagerty and Newsom, 2003), steam-heated systems (Swayze et al., 2014), Icelandic subglacial systems (Warner and Farmer, 2010) and low- temperature low-sulfur altered Icelandic basalts (Ehlmann et al., 2012), Nicaraguan fumaroles and mud pots (Hynek et al., 2013), to Yellowstone hot springs (Bishop et al., 2004), vent-proximal palagonitization on Mauna Kea, Hawaii (Golden et al., 1993), and acid-sulfate systems in

Svalbard, Norway (Hausrath, 2007; Treiman et al., 2002) and on Mauna Kea, Hawaii (Morris et al., 2005; Swayze et al., 2002). Compositional characterization is often drawn from a combination of instrumentation methods, including microprobe plus XRD (Hagerty and Newsom, 2003),

VSWIR plus XRD (Hynek et al., 2013; Swayze et al., 2002; Warner and Farmer, 2010), VSWIR,

XRD, SEM, and microprobe (Swayze et al., 2014), and VSWIR/Far-IR, XRD, x-ray fluorescence

(XRF), Mössbauer, electron microprobe, and electron microscopy (Golden et al., 1993; Morris et al., 2005). Many of these investigative techniques are not yet possible for Martian geological investigations. It is also important to note that for XRD analysis, the aforementioned work relied on additional sample preparation such as oriented mounts and glycolation (Hagerty and Newsom,

2003; Warner and Farmer, 2010). These methods enhance the ability to identify clay minerals with XRD (Środoń, 2006), but are not currently possible for in-situ analysis on Mars (Blake et al.,

2009).

A subset of these prior studies have also used Mars-relevant instruments when characterizing Mars-analog hydrothermal deposits (Bishop et al., 2004; Ehlmann et al., 2012;

Hynek et al., 2013). In a study of ten altered Icelandic basalt samples, Ehlmann et al. (2012) found that VSWIR reflectance spectroscopy did not provide a thorough assessment of the mineralogy, and would often not show zeolites which lack absorption features other than those related to water and other alteration products that could be identified via XRD. Conversely, VSWIR appeared to be the most useful tool for identifying various trace components such as smectite.

An analysis using VSWIR, mid-IR, XRD, thermal emission, and Raman spectroscopy of two hydrothermal deposits consisting primarily of calcite and siliceous sinter from Yellowstone

National Park by Bishop et al. (2004) found all methods to be successful at identifying calcite, but varying levels of identification for gypsum and aragonite. The amorphous and poorly crystalline

nature of the silicate deposits made identification with any one technique more difficult. A combination of methods resulted in the best possible characterization of this sample. Overall,

Bishop et al. (2004) found Raman to be most useful when identifying highly crystalline materials with polarized molecules, such as OH, whereas VSWIR was most useful for identifying sulfates,

OH bonds, and hydration.

Hynek et al. (2013) investigated samples from Telica, Momotombo, and Cerro Negro volcanoes in Nicaragua. All of the nearly 200 samples were analyzed with a field-portable Mars- equivalent XRD, and a subset of 29 samples were further analyzed with VSWIR. Overall, many of the same materials were detected by both instruments (SiO2, natroalunite, jarosite). However, many phyllosilicates and Fe-oxides were easier to identify in the VSWIR data. For phyllosilicates, this was attributed to the lack of sample processing before XRD analysis. The lack of Fe-oxides in the XRD data was attributed to their presence as a surface coating, which would be visible in the VSWIR, but not abundant enough to appear in the XRD bulk composition. Another explanation for the lack of Fe-oxides in the XRD data could be that they are present as nanocrystalline particles which is X-ray amorphous. Swayze et al. (2000) identified nanocrystalline goethite in a sample from an acid mine drainage site in Leadville, Colorado using

VSWIR, but when the goethite coating was removed and analyzed with XRD, no goethite was detected.

Icelandic field testing of an ExoMars PanCam emulator (AUPE-2) by Harris et al. (2015) demonstrated the ability to distinguish between various geologic units using a limited number of spectral bands between 400 and 1000 nm. The AUPE-2 was particularly useful for identifying variations in Fe content across visually indistinct soils. However, the limitations of a small spectral range and only 11 channels make specific mineral identifications rare. PanCam’s utility is the

ability to characterize outcrops in general terms, to be further investigated with other rover instrumentation.

The work presented here builds off the aforementioned studies – focusing on volcanic hydrothermal alteration in basaltic to basaltic andesite Mars analog systems, and using Mars- equivalent instrumentation (VNIR, XRD, and Raman) to best replicate the data that would be returned if these locations were remotely investigated using a rover payload. Analysis of a large sample set – containing 100 altered samples from seven different hydrothermal systems – provides further insight into potential hydrothermal products on Mars, and how these deposits may appear when investigated with in-situ instrumentation.

2.1.4. Field Sites

This study sampled across several volcanoes to ensure that a wide variety of alteration products were investigated – covering a range of geochemical conditions such as pH, temperature, and parent lithology. Sampling locations were chosen from previously-studied Mars analog sites, and included Poás (Rodríguez and van Bergen, 2017) and Turrialba (Beckerman, 2016) volcanoes in Costa Rica (Figure 4a), Cerro Negro, Momotombo, and Telica volcanoes (e.g. Hynek et al.,

2013; Marcucci et al., 2013; McCollom et al., 2013a) in Nicaragua (Figure 4b), and

Landmannalaugar and Krafla volcanoes (Hynek et al., 2014) in Iceland (Figure 4c). These seven sites encompass a wide variety of basalt to basaltic andesite compositions, all of which are comparable to compositions encountered on Mars or in the SNC meteorites (Table 4, Figure 5).

Sampling at Landmannalaugar included an altered rhyolite end member to represent alteration of the potentially evolved lavas on Mars (Wray et al., 2013).

Figure 4: Sampling locations in Costa Rica (Poás, Turrialba), Nicaragua (Momotombo, Telica, Cerro Negro), and Iceland (Krafla, Landmannalaugar). Stars represent capital cities.

Table 4: Terrestrial and Mars surface compositions of primary basalts, basaltic andesites, and rhyolites represented as oxide weight percents.

SiO2 Al2O3 FeOT MgO CaO Na2O K2O P2O5 Terrestrial sampling locations Cerro Negro, Nicaragua 1 49.73 19.50 9.70 4.73 11.50 2.18 0.43 0.12 Momotombo, Nicaragua 1 54.40 16.90 9.10 4.47 9.20 2.89 0.91 0.15 Telica, Nicaragua 1 51.72 19.10 9.60 4.13 10.10 2.84 0.96 0.17 Poás, Costa Rica 2 53.30 17.46 9.02 4.36 8.49 2.95 1.08 0.14 Turrialba, Costa Rica 2 54.14 16.30 8.46 6.02 9.22 3.12 1.43 0.37 Landmannalaugar, Iceland 3 69.04 15.03 3.10 0.34 1.76 5.82 4.48 0.03 Krafla, Iceland 4 48.54 15.39 11.58 8.80 12.71 1.83 0.11 n.r. Mars Adirondack, Gusev 5 45.30 10.43 21.09 11.90 7.76 2.09 0.03 0.54 Backstay, Gusev 6 49.40 13.10 13.10 8.30 6.00 4.00 1.02 1.34 Irvine, Gusev 6 47.50 8.30 19.70 9.50 5.80 3.00 0.60 0.94 Algonquin, Gusev 6 41.90 6.40 20.90 16.00 4.00 2.30 0.40 1.04 Bounce Rock, Meridiani 7 50.80 10.10 15.60 6.40 12.50 1.30 0.10 n.r. Shergottite LA 1 8 49.10 11.20 21.20 3.53 10.00 2.22 0.24 0.66 Shergottite QUE 94021 9 47.90 11.00 18.50 6.30 11.40 1.58 0.05 n.r.

All samples were analyzed with X-ray fluorescence. n.r.: Value not reported 1: Hynek et al. (2013) 2: This study 3: Blake (1984) 4: Nicholson and Latin (1992) 5: McSween et al. (2009) 6: Ming et al. (2008) 7: Rieder et al. (2004) 8: Rubin et al. (2000); Hurowitz et al. (2006) 9: Lodders (1998)

2.1.4.1. Poás Volcano, Costa Rica

Poás volcano is a basaltic to basaltic andesite composite volcano with extremely acidic (pH

 0), high temperature active fumaroles and crater lake. This dynamic geochemical system that pushes the temperature and pH boundaries is an excellent location to study alteration products in basalt-hosted lavas. Fresh Poás basalts collected in this study are primarily composed of plagioclase, with minor orthopyroxene and olivine, and are compositionally similar to Thermal

Emission Spectrometer (TES) measurements of the Mars surface (Table 4, Figure 5) (McSween et al., 2009). Although TES measurements are inherently surficial, and not likely to be representative of bulk rock composition (McSween et al., 2009), alteration products of analogous terrestrial basaltic andesites should still be studied to obtain a thorough understanding of all the potential secondary mineralogies we may encounter on Mars. Poás is particularly useful due to

the abundance of Fe-rich alunite in the alteration products, making it a useful location to study hydrothermal deposits containing intermediate members of the alunite-jarosite solid solution series. Identification of jarosite by the Opportunity rover’s Mössbauer spectrometer led to interpretations of an extremely low pH (<2) environment at some point in the geologic history of

Meridiani Planum (Morris et al., 2006; Squyres and Knoll, 2005). However, subsequent work by

McCollom et al. (2014) showed that jarosite and Fe-rich natroalunite are virtually indistinguishable using Mössbauer spectrometry, and further data, such as VNIR and/or XRD, are needed to distinguish between the two. The abundance of Fe-rich alunite/natroalunite in alteration products at Poás (up to 20% Fe substituted for Al, present in materials found outside of the hyperacid lake and not within the lake itself) suggests that these Mössbauer identifications of jarosite on Mars may actually be Fe-rich alunite/natroalunite, which would indicate a higher pH environment than jarosite, therefore changing our interpretation of the geochemical history of the region.

2.1.4.2. Turrialba Volcano, Costa Rica

Turrialba and Poás both have active fumaroles and crater lakes, but their chemistries, temperatures, and pH differ (Hynek et al., 2014; Prosser and Carr, 1987; Reagan and Gill, 1989).

Fresh Turrialba basalts gathered in this study are slightly more evolved than those at Cerro Negro, and are plagioclase-dominated, with clinopyroxene and olivine phenocrysts as well. Like Poás, basalts gathered from this study of Turrialba are compositionally similar to those measured by

TES (Table 4, Figure 5). The gas-dominated alteration regime at Turrialba has resulted in alteration products similar to mineral assemblages found on Mars, making it a useful analog environment for this study.

2.1.4.3. Cerro Negro Volcano, Nicaragua

Cerro Negro, a ~250 m high basaltic cinder cone, has experienced over 20 phases of eruptive activity since its formation in 1850 (Roggensack et al., 1997; Walker and Carr, 1986).

Sampling focused on alteration products from the 1992 and 1995 eruptive deposits – fresh basalts similar in composition to the SNC meteorites (Table 4, Figure 5) (Hynek et al., 2013, 2011), a commonly-utilized dataset for the estimation of Mars’ bulk composition (e.g., Taylor, 2013).

Petrographic analysis of fresh basalts from the 1992 and 1995 eruptions found plagioclase, augite, and minor forsterite phenocrysts (Hynek et al., 2013). Alteration at Cerro Negro is primarily due to the interaction of fumaroles and acid fog with the erupted basalts, and spans a wide range of temperature and pH values (Hynek et al., 2013).

2.1.4.4. Momotombo Volcano, Nicaragua

The Momotombo volcano is composed primarily of basalt and basaltic andesite (Table 4,

Figure 5), and has grown over the last 4500 years through several eruptive phases – over 15 of which have occurred in the last 400 years (Hynek et al., 2013; Novák, 2006). Fresh basalts at

Momotombo consist primarily of plagioclase and several types of pyroxene and are compositionally similar to TES measurements of the Martian surface (Hynek et al., 2013).

Alteration of this fresh basalt is through the interaction of very hot (up to ~450C) and highly acidic (pH < -1) fumarole emissions within the summit crater and on the flanks of the volcano.

2.1.4.5. Telica Volcano, Nicaragua

The Telica volcano is a highly active stratovolcano, with two large (~600 m diameter) summit craters. Highly active fumaroles within the craters and steep walls make access into the main craters difficult. Therefore, sampling at Telica was limited to materials on the crater rim and volcano flanks. Altered samples in these locations consist of rocks that were subjected to high temperature, low pH alteration at fumarole vents, and subsequently ejected from within the crater

during eruptions. Fresh materials at Telica are composed of basalt to basaltic andesite (Table 4,

Figure 5), and contain plagioclase, olivine, pyroxene, and quartz and are also compositionally comparable to TES measurements (Hynek et al., 2013).

2.1.4.6. Krafla Volcano, Iceland

The Krafla volcano straddles the rift zone in northeast Iceland, and includes a roughly 10- km-wide caldera, which is estimated to have formed roughly 100 kya (Saemundsson, 1991).

Deposits are primarily basalt lavas and hyaloclastites; however, there are areas that contain more evolved products such as dacite and rhyolite (Saemundsson, 1991). Krafla’s basalts are enriched in Fe (11.58 wt %) relative to most terrestrial basalts and similar in composition to basaltic shergottites (Table 4, Figure 5) (Nicholson and Latin, 1992). Samples from this site were collected from basalt-hosted fumaroles and hot springs at the Krafla caldera itself (Leirhnjúkur), as well as nearby Námafjall, and Þeistareykir.

2.1.4.7. Landmannalaugar Volcano, Iceland

The Landmannalaugar volcanic region lies within the Torfajökull central volcano in the central south volcanic region of Iceland. Unaltered substrate at this site consists primarily of rhyolite (Table 4, Figure 5) (Björke, 2010; Blake, 1984), in contrast to the other sample locations.

Although Martian surface composition measurements suggest global distribution of basalt to basaltic andesite (eg: McSween et al., 2009), Wray et al. (2013) identified three areas of highly evolved volcanics on Mars using CRISM data, granodioritic rocks have been identified in-situ in

Gale Crater (Meslin et al., 2013; Sautter et al., 2015; Williams et al., 2013), and felsic clasts have been identified in SNC meteorites (Humayun et al., 2013). These detections suggest areas of more felsic composition may exist both on Mars’ surface, and at depth (e.g. Horgan, 2013; Sautter et al.,

2016). Landmannalaugar serves as an analog site for these evolved Martian bedrock materials.

2.2. Methods

2.2.1. Sampling

A field campaign to Poás and Turrialba volcanoes was conducted in November 2013. 26 and 31 hand samples were gathered from within the Poás and Turrialba summit craters, respectively. An August 2012 expedition to Cerro Negro, Momotombo, and Telica volcanoes collected dozens of samples from each location. This study includes 5 representative samples from

Cerro Negro, 4 from Momotombo, and 5 from Telica, which were selected as representative samples out of a collection of nearly 200 and initially characterized by Hynek et al. (2013) using

VSWIR and XRD. Iceland sites – Landmannalaugar and Krafla (Leirhnjúkur, Námafjall, and

Þeistareykir) – were sampled in September 2013, with 9 and 21 returned representative samples, respectively. Sampling focused on areas with active fumaroles, areas of prior fumarolic activity, and areas in the current acidic steam cloud. In instances where the active region was not accessible, such as at Telica volcano, samples were gathered from the rim of the crater, and consisted of altered blocks that were ejected from the crater in previous eruptions. Sample collection included all visible mineralogical variations, based on color and texture, and was completed after a thorough walkabout examination of each site to ensure that the selected sample suite adequately reflected the mineralogical variability of the location.

2.2.2. Analysis

Analytical instrumentation was selected to mimic those that are available for rover-based investigations on Mars. In addition to operating as Mars-analogs, each investigative method has strengths that work well in conjunction with the others. VSWIR allows for identification of X-ray amorphous and Raman fluorescent materials such as phyllosilicates, which are key indicators of

geochemical conditions at the time of their formation. VSWIR also excels at identifying trace minerals, which may be present at levels below XRD detection thresholds. XRD is used for its ability to identify crystalline materials such as SiO2 polymorphs, sulfates, elemental sulfur, and highly crystalline Fe and Ti-oxides/hydroxides. Raman is also used in this study as a test of upcoming rover instrumentation, and for its sensitivity to various minerals such as elemental sulfur, anatase, sulfates, and Fe-bearing minerals. These three methods combined should allow for identification of the majority of mineral phases present.

2.2.2.1. VSWIR

The field-portable VSWIR spectrometer utilized in this study was the TerraSpec4 high resolution (Figure 6a) and the TerraSpec HALO reflectance spectrometers from ASD Inc. Both spectrometers have a 1 cm diameter sample area when used with a contact probe, and measure

2,151 channels from 350 to 2500 nm, with a 6 nm/band spectral resolution. The high signal-to- noise ratio allows for interpretation of relatively flat spectra, such as those that would be expected from fresh, unaltered basalts (ASD Inc., 2015a; Marcucci et al., 2013).

Spectra were analyzed using Excelis VIS’ Environment for Visualizing Images (ENVI) software, and absorption features were matched to USGS (Clark et al., 2013) and CRISM library spectra (PDS Geosciences Node and CRISM Spectral Library Working Group, 2014). Sample spectra were interpreted by matching them to library spectra of known compositions. It should be noted that library spectra are monomineralic, and are affected by grain size, water content, and slight chemical variations (Clark, 1999; Kirkland et al., 2003). However, sample spectra may contain mixtures of two or more materials, and/or have varying grain sizes and chemistries across small spatial scales. Hyperspectral VSWIR data was used for sample identification to simulate the use of a rover-based hyperspectral VSWIR instrument. Additionally, selected spectra were manually deconvolved to MER PanCam bands to investigate the utility of this dataset when working with complex mineralogical mixtures. Both regular and continuum-removed (baseline removed) sample spectra were compared to library reference spectra and analyzed for differences in absorption band depth due to changes in mineral abundance and grain size, and shifts in band location due to chemical variations.

2.2.2.2. XRD

Lab-based XRD analysis was done using the field-portable Terra XRD, which may be utilized both in the field and lab (Figure 6b). The Terra XRD is a functional-equivalent of the

CheMin XRD, but contains a Cu-K x-ray source, rather than cobalt. Cu-K x-rays result in compressed diffraction patterns – with smaller spacing between peaks – and an overall shift of the diffraction pattern to lower 2 values (Harris and White, 2007; Stahly, 2012). Although this source substitution results in slightly lower peak resolution, and the loss of some peaks at very low 2 values, a Cu-K source is considered to be an acceptable compromise for noise, signal attenuation,

and peak resolution (Harris and White, 2007) and comparable to the CheMin XRD for in-situ sample analysis (Blake et al., 2009; Sarrazin et al., 2005)

Samples were powdered to roughly 10 µm and a small amount (<< 1g) inserted into the mylar sample cell. To best mimic XRD analysis on Mars, no additional sample processing was completed. Samples were run for a minimum of 150 iterations to obtain the cleanest possible diffraction pattern. Resultant diffraction patterns were compared to the American Mineralogist

Crystal Structure Database “difdata” library diffraction patterns, supplemented with McCollom et al.’s (2014) experimental diffraction patterns of iron substitution in the alunite-jarosite solid solution series. All diffraction patterns were processed using XPowder software.

2.2.2.3. Raman

Laboratory Raman analysis was conducted using a Horiba LabRAM HR Evolution Raman

Microscope-Spectrometer (Figure 6c). This equipment utilizes both near-infrared (785 nm) and green (532 nm) lasers, and is capable of doing both single point analysis, and composition mapping across a field of view using a grid system for point analysis. This study focuses on analysis using several single point Raman spectra gathered throughout a sample, specifically targeted to identify all mineral phases present. Future work would benefit from incorporating high spatial resolution

(~ 2 m/pixel) Raman mineral mapping.

Raman analysis was done on whole rock, whenever possible, and powders when larger specimens were not available. Samples were shot with the 532 nm laser, which samples a wider spectral window (100 – 1700 cm-1), at various power levels and acquisition times. Initial spectra were corrected to remove baseline drift (ICS correction), and fluorescent effects removed.

Processed spectra were matched to library spectra from the RRuff database (Lafuente et al., 2015) using the KnowItAll software package. As there are currently no Raman laser spectrometers

operating on Mars for comparison, data collection and analysis followed standard terrestrial procedures.

2.3. Results

Altered samples were dominated by silica (amorphous and various crystalline phases),

Ca/Al/Fe/Mg-sulfates (gypsum, anhydrite, alunite, jarosite, hexahydrite, alunogen), and Fe-, and

Ti-oxides/hydroxides (hematite, goethite, anatase/brookite). Elemental sulfur and phyllosilicates

(montmorillonite, kaolinite) were also abundant (Table 5). Table 6 outlines key features for identification with each instrumentation method. Details regarding mineral identification will be discussed further in this section.

As anticipated, identification frequency varied by instrument (Table 7). Only definitive identifications (clear VSWIR absorption bands or XRD/Raman peaks that fit all the distinguishing features of a library spectra/diffraction pattern) were included in the count. Suggested presence

(very weak VSWIR absorption bands or small XRD/Raman peaks that required one or both of the other instrumentation methods used in this study for confirmation) was not counted as an identification. Each method was interpreted independently from the others to result in only those identifications that could be made with the individual dataset. Relative mineral abundance was not analyzed as part of this study; identifications were made solely as a present/not present basis.

Table 5: Abundance of mineral subgroups out of 100 Mars analog hydrothermal samples analyzed with VSWIR, XRD, and Raman. % of samples with positive identification (VSWIR, XRD, and/or Raman)

Amorphous SiO2 81% Crystalline SiO2 76% Cristobalite/Tridymite 63% Quartz 23% Sulfur 39% Sulfates 86% Gypsum/Anhydrite 44% Alunite-Jarosite series 63% Fe-sulfates 7% Mg-sulfates 33% Al-sulfates 22% Unidentified sulfates 5% Oxides/Hydroxides 94% Fe-oxides/hydroxides 76% Ti-oxides 77% Phyllosilicates 50% Montmorillonite 34% Saponite 5% Vermiculite 1% Unidentified smectite 5% Kaolinite 21% Unidentified zeolites 5% In instances where subgroup percents do not add up to major group percent, some subgroups may have been present in different samples (e.g. some samples contain both Fe- and Ti-oxides, while others contain just Ti-oxide).

Table 6: Characteristic features used for identification with VNIR, XRD, and Raman. VNIR (nm) XRD (2) Raman (cm-1)

SiO2 Amorphous Silica 1400 (as, d), 1900 (as), 2200 (as, d) 21.0 (bh) 500 (bh), 1100 (bh) Cristobalite No distinguishable features 22.0 (p) 112 (p), 230 (p), 418 (p) Tridymite No distinguishable features 20.5 (p), 21.6 (p), 23.2 (p), 29.95 (s), 207 (p), 280 (sh), 305 (p), 350 (p), 433 (p), 35.8/35.9/36.0 (t, s) 790 (s) Quartz No distinguishable features 20.9 (s), 26.7 (p) 110 (s), 132 (s), 209 (p), 357 (s), 405 (s), 466 (p) Sulfates Gypsum 1447/1491/1537 (t), 1747, 1944 (as), 1986 11.7 (p), 20.8 (p), 23.4 (p), 29.2 (p), 31.2 (p), 180 (s), 414 (s), 493 (s), 619 (s), 670 (s), 1008 (sh), 2177 (sh), 2216, 2265, 2438 33.4 (p) (p), 1135 (s) Anhydrite 632 (wk, sh); 1943 (wk) 25.5 (p) 498 (s), 608 (s), 675 (s), 1017 (p), 1130 (s) Alunite 1275 (wk), 1436/1495 (d), 1760, 1916 (wk), 15.5 (p), 17.9 (p), 25.5 (p), 29.9 (p), 31.0 (s), 206 (sh), 236 (p), 346 (s), 362 (s), 2168 (as), 2214 (wk, sh), 2322, 2400, 2446 39.4 (p), 47.8 (p), 52.4 (p) 465/485/510 (t, p), 560 (s), 655 (p), 1026 (p), 1079 (s) Jarosite 435 (wk), 639 (sh), 940 (br), 1469, 1521 (sh), 14.8 (p), 15.5 (p), 17.3 (p), 28.5/28.8 (d, p), 222 (p), 298 (s), 434 (p), 573 (s), 623 (p), 641 1849, 1935 (as), 2215/2265 (d), 2400 31.2 (s), 35.0 (s), 39.3 (s), 45.5 (s), 45.7 (s), (s), 1006 (p), 1100 (p), 1150 (p) 49.5 (s) Rhomboclase 428, 523,797, 1228 (wk, as), 1540 (br, as), 9.7 (p), 18.7 (s), 21.9 (s), 26.7 (s), 27.1 (s), 164 (s), 210 (sh), 240 (p), 264 (sh), 377 (p), 2005 (as) 28.6 (s) 449/470 (d, p), 615 (s), 1011/1030 (d, p), 1180 (p) Hexahydrite 434 (wk), 843 (wk, br), 1183 (wk, as), 1443 16.3 (p), 17.4 (p), 18.3 (p), 20.2 (p), 22.1 (p), 486 (s), 998 (p) (br, as), 1740 (br), 1945 (as), 2418 24.7/24.8 (d, s), 25.8 (s), 26.4 (s), 27.1 (s), 28.0 (s), 30.6 (p), 30.9 (p), 32.2 (s), 32.4 (s), 33.4 (s), 33.5 (s), 39.8 (s) Pickeringite 433, 883 (br), 1440 (br, as), 1764 (br), 1945 9.3 (p), 11.3 (s), 14.8 (s), 18.0 (p), 18.6 (p), 423 (s), 462 (s), 612 (s), 974/995 (d, p) (as) 19.1 (s), 19.4 (s), 20.7 (p), 21.4 (p), 21.7 (p), 22.4 (s), 22.6 (s), 23.7 (s), 25.5 (p), 25.9 (s), 26.9 (s), 28.2 (s), 30.2 (s), 31.0 (s), 31.7 (s), 32.5 (s), 33.0 (s), 33.5 (s), 34.5 (s) Alunogen ~1700, ~2200 6.6 (p), 19.8 (p), 20.3 (s), 20.6 (s), 21.8 (s), 422 (sh), 464 (s), 608 (s), 995 (p), 1082 (s), (no library spectra available) 23.0 (s), 24.2 (s) 1123 (s) Oxides/Hydroxides Hematite 537 (sh), 650 (sh), 744 (pk), 894 (br) 24.2 (p), 33.2 (p), 35.7 (p), 40.9 (p), 49.5 (p), 226 (p), 245 (s), 293 (p), 411 (p), 500 (s), 612 54.1 (p) (s), 660 (s), 1320 (bh) Goethite 494 (sh), 667 (sh), 767 (pk), 929 (br) 17.8 (s), 21.3 (p), 33.3 (p), 34.7 (s), 36.0 (s), 246 (p), 300 (p), 387 (p), 400 (s), 417 (sh), 481 36.7 (p), 40.0 (s), 41.2 (s), 53.3 (p) (p), 547 (s), 1000 (s) Anatase No distinguishable features 25.3 (p), 36.8 (s), 37.7 (p), 38.5 (s), 48.0 (p), 145 (p), 394 (s), 514 (s), 637 (s) 53.7 (p) Phyllosilicates Montmorillonite 1412, 1466 (sh), 1906, 1960 (sh), 2204 6.4 (p), 19.6 (p), 21.5 (s), 34.8 (s) Fluorescent Saponite 1389/1410/1463 (t), 1779 (sh), 1905, 1972 Mineral identification requires additional Fluorescent (sh), 2289 (sh), 2313, 2386 sample processing Nontronite 460 (sh), 518 (sh), 653, 965 (br), 1413/1434 6.2 (p), 19.5 (s) Typically fluorescent, but can have peaks at (d), 1908 (as), 1981 (sh), 2287, 2398 179 (p), 233 (p), 283 (p), 357 (s), 424 (p), 510 (p), 583 (s), 678 (p), 811 (s), 883 (p), 970 (s), 1033 (p) Kaolinite 1395/1414 (d), 1911, 2164/2205 (d), 2289 12.4 (p), 20.3 (s), 21.2 (s), 23.2 (s), 24.9 (s) Fluorescent (wk), 2316 (wk), 2354 (wk), 2381 (wk) Zeolites Laumontite 981, 1165, 1431, 1789 (br, as), 1923 (as) 9.3 (p), 12.9 (p), 19.7 (s), 21.4 (p), 25.3 (p) 198 (s), 309 (p), 326 (sh), 382 (s), 492/512 (d, p), 590 (s), 671 (s) Other Elemental Sulfur 419 (sh) 11.5 (s), 15.4 (s), 21.9 (s), 22.7 (s), 23.1 (p), 90 (s), 154 (p), 220 (p), 245 (s), 437 (s), 472 25.9 (p), 26.7 (p), 27.8 (p), 28.7 (p), 29.0 (p), (p) 31.4 (s), 34.2 (s), 35.9 (s), 37.1 (s), 37.4 (s), 38.0 (3), 39.4 (s), 42.8 (s) Continuum-removed VSWIR absorption features are described as: as = asymmetrical; sh = shoulder; br = broad; wk = weak; d = doublet; t = triplet, pk = peak. XRD and Raman peaks are described as: p = primary; s = secondary; bh = broad hump; sh = shoulder; d = doublet; t = triplet; Raman fluorescence interferes with mineral identification; VSWIR features from Clark et al. (2013) and PDS Geosciences Node and CRISM Spectral Library Working Group (2014), XRD and Raman features from Lafuente et al. (2015)

Identified when present VNIR XRD Raman

Amorphous SiO2 81% 83% 23% Crystalline SiO2 9% 95% 9% Elemental Sulfur 59% 56% 69% Sulfates 38% 57% 45% Oxides/Hydroxides 26% 55% 56% Phyllosilicates 91% 24% 0% Zeolites 57% 29% 14%

SiO2 Amorphous Silica 81% 83% 23% Cristobalite/Tridymite 0% 100% 6% Quartz 9% 96% 22% Sulfates Gypsum/Anhydrite 25% 80% 50% Alunite-Jarosite series 56% 90% 65% Fe-sulfates 29% 57% 29% Mg-sulfates 85% 18% 21% Al-sulfates 77% 82% 23% Unidentified sulfates 40% 20% 40% Oxides/Hydroxides Fe-oxides/hydroxides 61% 42% 75% Ti-oxides 3% 53% 100% Phyllosilicates Montmorillonite 88% 24% 0% Saponite 100% 0% 0% Vermiculite 0% 33% 0% Unidentified smectite 100% 0% 0% Kaolinite 76% 71% 0% Zeolites Unidentified zeolite 100% 67% 33% Other Elemental Sulfur 59% 56% 69%

2.3.1. SiO2

Amorphous SiO2 was most readily identified via VSWIR and XRD (Figure 7).

Identification with VSWIR focused on the location and shape of the ~1400 and ~1900 nm H2O/OH overtones and combinations, as well as the ~2200 nm Si-OH combination stretch and bend. The

2200/1900 nm and 1410/1460 nm band depth ratios from Rice et al. (2013) were used to confirm the presence of opal-A versus more crystalline opal-CT, opal-C, and microcrystalline quartz.

Amorphous SiO2 was also detectable in XRD patterns as a broad hump centered around 22 2.

Although the presence of amorphous SiO2 may also be suggested through broad Raman peaks around 1100 and 1600 cm-1, this was not found to be a diagnostic method for these materials as

they were often indistinct, and data acquisition challenges were common due to high levels of fluorescence in these samples.

Crystalline forms of SiO2 (cristobalite, tridymite, and quartz) were almost exclusively identified via XRD (Figure 7). Sharp, and often large, diffraction peaks at ~22 2 (cristobalite),

20.5, 21.65, 23.2 2 (tridymite), and 20.9, 26.67 2 (quartz) allowed for unequivocal identification of these crystalline SiO2 forms. Cristobalite and quartz were also identified through

Raman peaks at 112, 230, and 419 cm-1 (cristobalite), and 210 and 468 cm-1 (quartz) (Figure 7).

Conversely, these minerals could not be identified through VSWIR due to their relatively featureless spectra throughout the 350 – 2500 nm range; and although tridymite was found in 9 samples using XRD, no tridymite was detected using Raman spectroscopy, though it should be possible to identify via multiple peaks at 210, 280, 305, 350, and 433 cm-1.

2.3.2. Elemental Sulfur

Elemental sulfur was most reliably identified with Raman, although it often produces distinct peaks in XRD patterns, and an indistinct VSWIR shoulder as well (Figure 8). When in contact with the Raman laser, sulfur creates extremely large peaks (90, 155, 220, 245, and 470 cm-

1) relative to the baseline signal. A major challenge with Raman sulfur identifications is the ability of the sulfur signal to easily overpower that of any secondary materials that are present in the sample.

Figure 8: Details of elemental sulfur identifications in VSWIR, XRD, and Raman using representative examples – easily identifiable using all three methods. In panel B, VNIR spectra are displayed as continuum-removed for more accurate band center locations. The red VSWIR curve is the USGS splib06 library reference spectra for sulfur, blue is the USGS splibo6 library spectra for K-alunite, and pink is the McCollom et al., (2014) Fe-rich (10%) natroalunite library spectra. VSWIR spectra a and b (CR13 Turrialba 3 and I13 P0 yl) display the characteristic sulfur shoulder, while spectra c (CR13 Poás Hot 5) does not, despite containing ~ 19 wt% sulfur via XRD analysis.

Sulfur’s highly crystalline nature also allows it to be easily identified with XRD. Several sharp diffraction peaks may be present, depending on the crystal form in the sample. The most common sulfur diffraction pattern found in our samples had major peaks at 23.1, 25.8, and 27.7

2 and several minor peaks listed in table 4. Unlike in Raman data where sulfur overpowers other materials, sulfur is easily identified in XRD without masking other mineralogical components.

Although it is possible to identify elemental sulfur with VSWIR, this method is less reliable than the aforementioned Raman and XRD. VSWIR identifications of sulfur rely on the ~450 nm sulfur conduction band, which creates a large shoulder around 450-500 nm. This shoulder is often lost in complex mineral mixtures (Figure 8) – around ~ 19 wt % when mixed with Mg- or Fe- sulfates (hexahydrite/pickeringite, jarosite, or Fe-rich (natro)alunite), hematite, and/or saponite in our samples – and is not a reliable diagnostic method of sulfur in these mixtures.

2.3.3. Sulfates

Sulfates were most easily identified with XRD (identified with XRD in 58% of samples containing sulfates), although the identification rates for this mineral class were also 38% for

VSWIR and 45% for Raman (Table 7). Common hydrothermal sulfates, such as gypsum (CaSO4

* 2 H2O), alunite (KAl3(SO4)2(OH)6), jarosite (KFe3(SO4)2(OH)6), hexahydrite (MgSO4 * 6 H2O), pickeringite (MgAl2(SO4)4 * 22 H2O), apjohnite (MnAl2(SO4)4 * 22 H2O), and alunogen

(Al(SO4)3) have highly ordered crystal structures that result in XRD patterns and Raman spectra with several distinct, large peaks, and are easily identifiable, even in complex mixtures (Table 6,

Figure 9).

Several specific VSWIR absorption bands can be attributed to sulfate minerals (Table 6,

Figure 9). All sulfates in this study exhibit the ~2400 nm OH/H2O + /OH/H2O combination (Cull-

Hearth et al., 2016), as well as the ~1770 nm OH bend and stretch combination and overtone. The

~1400 and ~1900 nm H2O/OH combination and overtone bands are often present, and vary depending on hydration state. Fe-bearing sulfates also displayed absorption bands due to Fe spin transitions between 400 – 900 nm (Fe3+) and 900 – 1200 nm (Fe2+) (Cull-Hearth et al., 2016).

Various metal-OH bands between 2000-2500 nm can be attributed to Al-OH (~2200 nm), Fe-OH

(~2900 nm), and Mg-OH (~2300 nm), and also help distinguish between sulfate minerals.

2.3.4. Oxides and Hydroxides

Raman and XRD exceled for crystalline Fe- and Ti-oxide/hydroxide identification; however, Fe-oxides/hydroxides were also identifiable using VSWIR. Although VSWIR was more reliable than XRD for Fe-oxide/hydroxide identification, this method does not appear to be as reliable as Raman, which is sensitive to Fe-bearing minerals (Table 7) due to their tendency to burn easily when in contact with the laser. Hematite (Fe2O3) and goethite (FeO(OH)) were present as weathering rinds and may have been in a nanocrystalline form on many samples, which may contribute to the lower identification rates for XRD, and will be addressed in the discussion section of this paper.

Hematite and goethite were easily identified through all three instrumentation methods

(Figure 10). We used Fe3+ electronic transitions at ~640 and 900 nm for both goethite and hematite, and ~430 and 550 nm for hematite to identify these materials in VSWIR (Brown et al.,

2008). For XRD, hematite was identified via peaks at 24.2, 33.2, 35.8, 41.0, 49.5, and 54.2 2, and goethite via peaks at 21.3, 33.2, and 36.6 2. In Raman spectra, a broad hump around 1300-

1350 cm-1 with peaks at 120, 225, 245, 300, and 415 cm-1 (Lafuente et al., 2015) were indicative of hematite. For goethite, the broad hump shifts to ~1300 cm-1, and peaks are located at 297, 397,

550, and 687 cm-1 (Lafuente et al., 2015). The challenges of Fe-bearing mineral identification with Raman spectroscopy lie in the tendency of Fe-bearing materials to burn easily when in contact with the laser. However, this characteristic burning can also be helpful for mineral identification, and does not preclude the collection of distinct Raman spectra for these minerals.

Ti-oxide polymorphs anatase, brookite, and rutile (TiO2) were all present in our samples, and most readily identified through Raman spectroscopy. Similar to elemental sulfur, TiO2 results in intense Raman peaks (140, 395, 520, and 638 cm-1 for anatase – the most common polymorph

in our samples) and can easily overwhelm the signal from any other minerals present in a sample.

Although TiO2 is also easily identifiable in XRD, with distinct peaks at 25.3, 37.8, and 48 2, the signal is less pronounced and does not overpower those from other materials. VSWIR is not a useful tool for TiO2 identification due to the location of the main diagnostic absorption band at

~340 nm – just outside of the VSWIR window.

2.3.5. Zeolites

Zeolites were tentatively detectable in 5 of our samples using VSWIR and XRD.

Laumontite (CaAl2Si4O12 * 4 H2O) was identified in XRD by large peaks around 9.5, 13, and 21.5

2. Laumontite identification in VSWIR is primarily through the location and shape of the ~1420 and ~1920 H2O/OH combinations and overtones, however, identifying zeolites through only

OH/H2O absorptions is unreliable due to their presence in many water-bearing minerals. These identifications should be confirmed through either Raman or XRD. Although it is possible to identify zeolites with Raman, we were not able to confirm the presence of zeolites using Raman in any of the 5 samples with VSWIR or XRD detections. The low number of zeolite detections (5 possible out of 100 samples), nearly equal identification rates for VSWIR (3 of 5) and XRD (2 of

5), and identification discrepancies between detection methods do not suggest either method is superior for zeolite identification.

2.3.6. Phyllosilicates

As expected, VSWIR out performed XRD and Raman for phyllosilicate identification, but

XRD was also able to identify some montmorillonite (Na,Ca)0.33(Al,Mg)2(Si4O10)(OH)2 * n H2O), vermiculite ((Mg,Fe)3[(Al,Si)4O10](OH)2 * 4 H2O), and a large percentage of the kaolinite

(Al2Si2O5(OH)4) present (Figure 11). Due to the highly fluorescent and poorly crystalline nature of phyllosilicates, Raman spectroscopy was not able to identify these materials.

Montmorillonite, saponite (Ca0.25(Mg,Fe)3((Si,Al)4O10)(OH)2 * n H2O), nontronite

((CaO0.5,Na)0.3Fe2(Si,Al)4O10(OH)2 * n H2O), vermiculite, kaolinite, and a general (unclassified) smectite could be identified in VSWIR by variations in the ~1400 and ~1900 nm H2O/OH combinations and overtones – although the ~1400 nm bands were often masked by the presence of amorphous SiO2 and were not as reliable an indicator as others. Fe-bearing phyllosilicates also

displayed Fe electronic transitions between 400 and ~1000 nm. The primary diagnostic bands for phyllosilicate identification were the metal-OH combinations and overtones from 2100 to 2500 nm.

Identification of phyllosilicates in XRD, while possible, is not as effective as VSWIR.

Although montmorillonite, vermiculite, and kaolinite do have distinguishable peaks in XRD, these are highly dependent on sample preparation and relative abundance. Other phyllosilicate distinctions are not possible in this study due to their reliance on additional sample processing, which was not conducted in an effort to mimic current sampling methodology on Mars.

Montmorillonite was identifiable in our samples due to a single small peak at 6 2, vermiculite with a single large peak at ~6.5 2, and kaolinite with a small-to-moderate peak at ~12.3 2.

Relatively kaolinite-rich samples such as I13 KL cs min2 (an ~18 wt % kaolinite sample from

Krafla, Iceland) also displayed secondary peaks noted in table 4 (Figure 11d). However, most samples only contained the ~12.3 2 kaolinite peak. General indications of phyllosilicates were also possible when the diffraction pattern contained a large negative slope at low 2 values, and a broad hump within that negative slope, as seen in the sample CR13 Poás Lake 4, a montmorillonite and kaolinite-bearing sample from the banks of Poás’ acidic crater lake (Figure 11d). To identify a specific phyllosilicate mineral in these samples with XRD, further sample preparation would be needed.

2.4. Discussion

2.4.1. VSWIR

The surficial nature of VSWIR sampling, nonlinear mixing, and masking effects (eg: darker color minerals dominating the reflectance spectra) may all limit VSWIR’s identification capabilities (e.g. Clark et al., 1990). With these materials this method cannot be used to identify quartz, cristobalite, or anhydrite when they are mixed with other materials, due to their relatively featureless spectra in the VSWIR range. VSWIR is most useful when it comes to identifying phyllosilicates (vs Raman or XRD) and equally effective as XRD for identification of amorphous silica, but less so for sulfates. This contrasts with Bishop et al.’s (2004) results suggesting VSWIR is the best method for sulfate identification, and may be due to coatings of other minerals such as phyllosilicates or Fe-oxides masking the VSWIR signal from sulfates. Bishop et al. (2004) investigated two samples containing only one sulfate, gypsum, which provides a limited dataset for comparing sulfate identification methods. Additionally, for sample analysis, both surface coatings and interior powders were analyzed, which may have allowed for identification of gypsum even with a surface coating. Perhaps surprisingly, identification rates for elemental sulfur and Fe-oxides/hydroxides were slightly higher with VSWIR than XRD. This may be attributed to the presence of surface weathering rinds that are identifiable with VSWIR, but not abundant enough in the bulk composition to appear in the XRD pattern, similar to the results from Hynek et al. (2013) and Marcucci et al. (2013). Alternatively, this may be due to trace amounts of these materials throughout the sample, below the ~3 wt. % necessary for identification with XRD

(Sarrazin et al., 2005).

Overall, VSWIR proves to be a useful tool due to its sensitivity to minerals at low relative abundance and may be used to identify bulk mineralogy at a variety of spatial scales. However,

VSWIR is difficult to use for quantitative analysis, and may miss bulk composition components due to the surficial nature of reflectance spectroscopy as well as masking effects of weathering rinds of mixed compositions. Additionally, reference library spectra for common hydrothermal products (eg: cristobalite, alunogen) are not included in commonly utilized publicly-available reference libraries such as USGS and CRISM and are difficult to find. This lack of reference spectra is not a failure of instrumentation, but may still result in interpretive errors, especially if algorithms are used for spectral interpretation. Because of this, the VSWIR science community would benefit greatly from a concentrated effort to fill the gaps in publicly available reference libraries.

2.4.2. XRD

XRD is the standard method for thorough bulk identification of sample mineralogy. Our results show XRD’s strengths lie in identification of highly ordered crystalline sulfates, sulfur, and silica phases (both crystalline and amorphous). The inability of XRD to identify sulfates in many of the samples where these materials were found with VSWIR and/or Raman may be due to their presence as trace minerals, and therefore below the XRD detection threshold. Although XRD was also capable of identifying Fe- and Ti-oxides/hydroxides, they were more readily identified with

Raman (Ti and Fe) or VSWIR (Fe). As noted above, this is likely to be due to their presence in thin weathering rinds.

In x-ray diffraction patterns, coarser grained materials result in high relative peak intensities, and can often mask the lower relative intensity peaks of fine-grained materials

(McCusker et al., 1999; Środoń, 2006) such as phyllosilicates – common hydrothermal products which are excellent indicators of geochemical conditions. Our results show nearly equal frequency of kaolinite identification in XRD (71%) as VSWIR (76%) when present in a sample, but

significantly lower identification rates for other phyllosilicates such as montmorillonite (24% and

88% for XRD and VSWIR, respectively). If kaolinite is present in trace amounts (< 3 wt % for the Terra XRD utilized in this study), it is likely that XRD will not be able to identify it without additional sample processing, while VSWIR will still be able to detect this mineral phase. The higher identification rates for kaolinite may be attributed to the large amounts of kaolinite produced by argillic alteration (Heald et al., 1987; Pirajno, 2010) – thereby resulting in samples that have a high relative abundance (> 5 wt %) of kaolinite, which is able to be identified without the use of additional sample processing methods. Additionally, kaolinite’s distinction as a coarse- grained phyllosilicate may also make it the easiest phyllosilicate to identify in mixed samples

(Środoń, 2006), as the nature of XRD results in strong peak signals from larger grained materials.

Kaolinite also displays several primary diffraction peaks that range from 12.4 to 24.9  2, and multiple secondary peaks (Table 6), making them easier to identify compared to other phyllosilicates with diagnostic peaks that are limited to lower 2 values and harder to distinguish.

In spite of this, accurate phyllosilicate identification for samples containing < 3 wt % phyllosilicates requires an intermediate processing step which is not currently possible for in-situ rover analysis (Beckerman, 2016; Bodine and Fernalld, 1973; Vaniman et al., 2014). Without this intermediate processing step, phyllosilicate presence may be suggested, but VSWIR is needed for confirmation. In spite of this, phyllosilicate identification at Gale crater relies heavily on XRD

(eg: Vaniman et al., 2014) in the absence of rover VSWIR capability, and may be leading to scientific interpretations that are missing key indicators of environmental conditions.

2.4.3. Raman

Raman spectra gathered in this study are clear and easy to interpret. However, elemental sulfur’s high level of Raman reactivity often masks the signal from other minerals, and is

problematic in sulfur-rich areas such as hydrothermal deposits. Additionally, Fe-rich materials burn easily (and may therefore result in corrupted spectra) and phyllosilicates and amorphous materials fluoresce strongly when in contact with the laser. It is possible to gather clean Fe-rich spectra, but extreme care must be exercised. Our results suggest Raman laser spectrometry is most effective for the identification of elemental sulfur, Fe-oxides/hydroxides, and Ti-oxides due to their high Raman activity. Raman is also useful for the identification of sulfates, but less so than

VSWIR and XRD. This is likely due to the high Raman reactivity levels of elemental sulfur and

Ti-oxides masking the signals those of other less abundant and/or Raman active minerals, and the tendency for materials coated with Fe-rich minerals to burn before a spectra can be gathered.

Raman spectrometers that utilize a 532 nm laser may also be used to suggest the presence of amorphous or poorly crystalline phases (silica, clays) through fluorescence and broad peaks in the spectra, but additional methods are needed for confirmation. The Mars 2020 SHERLOC Raman uses a DUV laser, which minimizes fluorescence and should be capable of identifying various minerals, including phyllosilicates (Beegle et al., 2015).

Raman’s point analysis technique results in clean and easy to interpret spectra. However, point analysis is limited to a very small area (~50 m for rover-based Raman), and extensive investigation of a sample must be done to find all the present mineral phases. It is easy to miss less abundant phases with this method, and this may have contributed to some of the non- identifications with Raman in this study. One potential solution to this problem is to utilize mineral mapping via high resolution grid analysis to cover a larger area and increase the chances of encountering all the mineral phases present in a sample. High resolution grid mapping is a common feature of terrestrial lab-based Raman spectrometers, and an included feature for both the

ExoMars and Mars 2020 Raman Laser Spectrometers (Table 3). Micro-mapping mode for the

Mars 2020 SHERLOC Raman will collect 400 laser pulses for each 50 m point location. Each set of 400 laser pulses will take 10 seconds to complete. With 400 individual point locations to sample, this dataset will take just over an hour to gather (Beegle et al., 2015). While providing a wealth of mineralogical information, grid mapping may become difficult for samples lacking flat surfaces. Irregular surfaces can result in areas that are out of focus, with poor data quality. The rover teams could circumvent this by analyzing the smooth sides of drill cores or freshly drilled surfaces; however, it will remain to be seen whether these surfaces will be suitably flat for detailed mineral mapping. ExoMars’ RLS is equipped with a 1 mm autofocusing lens (Maurice et al.,

2013), and Mars 2020’s SHERLOC DUV Raman also contains an autofocusing lens with a range of 12.5 mm (Beegle et al., 2015). Both of these should help counteract any rough surfaces. One additional challenge for mineral mapping is deposits that are rich in amorphous or poorly crystalline phases, such as amorphous silica or phyllosilicates. The high levels of fluorescence in these samples (for Raman using 532 nm lasers) makes it difficult to determine an appropriate laser setting to be used over the whole mapping area. For both of these challenges, a thorough point analysis is most effective.

2.4.4. Implications for Mars

2.4.4.1. MER Spirit at Gusev Crater

Of the investigative methods utilized in this study, the MER Spirit rover carries only one

– the PanCam broad-band VNIR imager. Additional compositional analyses must be done by

APXS, Mössbauer, and Mini-TES. While each of these instruments has their strengths, APXS results in elemental abundances and must be extrapolated to hint at the mineralogical composition of a sample (Rieder et al., 2003). Mössbauer, while a useful tool for investigating mineralogy, is specifically designed for the identification of Fe-bearing minerals (Morris et al., 2006; Wdowiak

et al., 2003), and is less-well suited for other materials, such as elemental sulfur, SiO2 polymorphs, and Al- and Mg-rich phyllosilicates or sulfates – all of which are common materials in hydrothermal systems such as Home Plate in Gusev Crater. Mini-TES is also available as a mineralogical tool, but like PanCam, is sensitive to surficial deposits such as weathering rinds and dust (Christensen et al., 2003) – although the contribution of dust can be somewhat accounted for using an endmember spectra (Christensen et al., 2004).

The MER PanCam samples only 11 spectral bands – the same spectral bands utilized on

Pathfinder (Bell et al., 2003) – from 400 to 1000 nm, resulting in a very limited dataset. PanCam is primarily used for context imaging and is best suited for identification of Fe-bearing minerals such as hematite, goethite, and Martian dust by analyzing absorption shoulders from various Fe electronic transitions. PanCam can also give some insight into the hydration state of minerals due to the ~950 nm OH/H2O combination and overtone. However, the narrow spectral window misses other critical absorptions between 1000 and 2500 nm – key indicators of hydration states, zeolites, sulfates, opals, and phyllosilicates. Rather than providing detailed mineralogical identification,

PanCam is often utilized to create false color images (R: 753 nm; G: 535 nm; B: 432 nm), which are used to observe surficial variations and identify targets for more detailed investigation with other instruments.

Wang et al. (2006c) suggested the presence of 14 – 17 wt % kaolinite-serpentinite group phyllosilicates (kaolinite, and potentially antigorite and greenalite) at the Wooly Patch outcrop on the West Spur of Columbia Hills. Cation ratios, CWIP normative plots, and mineral mixing models were based on an abundance (relative to the previously analyzed plains basalts) of Al2O3

2+ and SiO2 in the elemental data from APXS, Fe coordination from Mössbauer, as well as a strong

Fe absorption band in the PanCam broad-band VNIR. However, no direct mineral detections were

made due to the limited spectral range of the PanCam VNIR, and extremely cold temperatures rendering the mini-TES data uninterpretable. Although a robust analysis was completed, Wang et al. (2006c) expressly state the need for direct mineralogical detections to confirm their results.

These could have been possible if the spectral range had covered the diagnostic 2100/2200 nm Al-

OH doublet exhibited by kaolinite. Laboratory VSWIR spectra of samples collected for this study that display strong kaolinite absorption bands were convolved to MER PanCam band wavelengths

(Figure 12). After this convolution, only one kaolinite-bearing sample was potentially identifiable

(I13 K Viti), due to the lack of Fe-oxides/hydroxides in this sample. All other kaolinite-bearing samples had strong absorption features from other minerals (primarily Fe-oxides/hydroxides) masking the relatively featureless kaolinite spectra at short wavelengths (400 – 1100 nm).

Although kaolinite is difficult to detect using the MER PanCam, given our results (Figure 11d), the 14-17 wt % kaolinite should be high enough to detect through XRD even without additional sample processing.

Figure 12: Laboratory spectra of phyllosilicate-bearing samples collected for this study, manually deconvolved to MER/MSL PanCam bands (panel A). All of the collected samples shown here contain kaolinite, as evidenced by the strong ~2100/2200 nm doublet in panel C (continuum- removed spectra for more accurate location of band centers). Nontronite (I13 KL red and I13 Nama U rd 1) and potentially saponite (CR13 Poás Lake 3) are also present in various samples, and may be identified by the ~2286 nm (nontronite) and ~2315 nm (saponite) absorption bands, also shown in panel C. Using MER/MSL PanCam band locations (panel A), kaolinite is not identifiable in any of these samples due to the lack of characteristic absorption bands at shorter wavelengths, and nontronite would be incorrectly identified as being present in both the CR13 Poás Lake 3 and CR13 Poás Lake 4 samples. At shorter wavelengths, the VNIR spectra appear to be dominated by Fe-oxides – the I13 KL red and I13 Nama U rd 1 samples both display the 532 and 834 nm Fe electronic transition absorptions indicative of hematite (panels A and B), which mask the 934 nm absorption feature that may indicate the presence of nontronite, but may also be due to other minerals such as Fe-bearing montmorillonite, or the Fe-hydroxide, lepidocrocite.

In their analysis, Wang et al. (2006c) assume all of the abundant TiO2 (0.94 and 0.91 wt % for the “Sabre” and “Mastodon” targets, respectively) is present in ilmenite – a primary basaltic accessory mineral. However, anatase (TiO2) is an extremely common product of hydrothermal alteration (present in 77% of our collected samples), and should be considered as a likely source of TiO2 in the Wooly Patch outcrop. Although unidentifiable using VNIR, anatase could be easily identified with either XRD or Raman, and would further support Wang et al.’s (2006c) proposed acid leaching scenario for the formation of Wooly Patch.

In addition to phyllosilicates, an abundance of opaline silica has been identified at Home plate through Spirit’s PanCam instrument via an absorption at the 1009 nm wavelength (Rice et al., 2010; Ruff et al., 2011) as well as Mini-TES, APXS, and Microscopic Imager (MI) observations (Ruff et al., 2011). The 1009 nm PanCam absorption serves as an indicator that opaline silica is present; however, as Rice et al. (2013) noted the ~950 nm H2O absorption band that is causing this downturn at the end of the PanCam spectra is located between two of the

PanCam filter wavelengths (934 and 1009 nm), and is therefore not detectable in all silica- containing samples. Only those with the strongest H2O features will be detectable using PanCam.

VSWIR spectra of several of our amorphous SiO2-bearing samples (identified in both VSWIR and

XRD) were convolved to MER PanCam VNIR bands (Figure 13), and of the 10 example spectra, only two (CR13 Turritop1 and CR13 Turrialba N7 – two hydrated SiO2-bearing samples from within the Turrialba summit crater) display the 1009 nm absorption band. The remaining 8 examples would not have been identified as containing SiO2. This suggests many SiO2-bearing deposits in Gusev may have gone unidentified. In addition to the potential to miss the ~950 nm

H2O absorption band, the key features for opal classification – and subsequent insight into geochemical conditions of formation – are the 1410, 1900, 2210, and 2260 nm absorption bands

(Rice et al., 2013), which are not included in PanCam’s range. Ruff et al. (2011) presented Mini-

TES data to suggest the presence of opal-A rather than more ordered phases; however, due to dust contamination, viewing angle effects, and noise this identification remains speculative. Without the longer-range VSWIR absorption features from 1400 – 2260 nm, the Spirit rover data was not useful for distinguishing between opal-A and opal-CT – a distinction which would provide a great deal of insight into temperature and the availability of water at the time of formation, and therefore the potential habitability of Home Plate (Rice et al., 2013). From our results, it is clear that a lack of mineral signature in PanCam data should not cause science teams to conclude that the mineral is not present in a deposit. Further analysis via the additional onboard instrumentation must be done to provide more thorough insight into the complete mineralogical assemblage for any given outcrop. Additionally, relying heavily on PanCam data for science target selection will likely cause investigative teams to miss key minerals, and further site selection should consider that geochemically-relevant minerals such as phyllosilicates and hydrated SiO2 may not be accurately represented.

Figure 13: Laboratory spectra of amorphous SiO2-bearing samples collected for this study, deconvolved to MER/MSL PanCam bands (panel A). All of the collected samples shown here contain opal or siliceous sinters, as evidenced by the strong ~1410/1460 nm doublet in panel C, asymmetrical ~1900 nm band in panel D, and ~2210/2260 nm doublet in panel E (continuum- removed spectra for more accurate location of band centers). Presence of amorphous SiO2 is also confirmed via XRD for each sample, and distinctions between opal-A, opal-CT, and sinters were made using VNIR library spectra as well as 1410, 1460, 1900, and 2200 nm band depth and ratios as described by Rice et al. (2013). Although amorphous SiO2 is present, it is often not detectable with MER/MSL PanCam bands – as evidenced by the lack of downturn at 1009 nm. Often this lack of a 1009 nm absorption is due to the presence of hematite, goethite, or other Fe-bearing materials, which dominate the VNIR spectra at shorter wavelengths, as seen in samples CR13 Poás Fumarole 12 and I13 Nama U or1, I13 Nama Core1, CR13 Poás Fumarole 11, and CR13 Turritop 1. Of this subset, the only samples that would appear to contain amorphous SiO2 at MER/MSL resolution are CR13 Turritop1 and CR13 Turrialba N7. Hydrated SiO2 in all other samples would likely not be identified.

2.4.4.2. MER Opportunity at the Matijevic formation, Endeavour crater

The MER Opportunity rover recently reached the Matijevic formation on the east side of

Cape York – part of Endeavour crater’s rim. This outcrop was of particular interest due to a nontronite/saponite-like spectra observed in a CRISM footprint of the Cape York region displaying absorption bands around 2300 and 2400 nm (Arvidson et al., 2014; Wray et al., 2009). This phyllosilicate outcrop may be attributed to pre-impact acidic alteration (Arvidson et al., 2014;

Wray et al., 2009) or impact-induced hydrothermal alteration (Schröder and Schwenzer, 2017).

The loss of Opportunity’s Mössbauer and Mini-TES instruments have limited investigation of the Matijevic formation to PanCam and APXS. In-situ PanCam observations of RAT-brushed

Matijevic targets show a weak absorption at the 932 nm band location and a deeper (relative to the

Matijevic matrix) 535 nm absorption from the dark surface coatings, which may be attributed to higher levels of crystalline Fe-oxides (532 nm) and an Fe-bearing smectite such as nontronite (932 nm) (Farrand et al., 2014). While this ~932 nm absorption band could potentially be attributed to the ~955 nm nontronite Fe transition, Farrand et al. (2014) note that several minerals, such as Fe- bearing montmorillonite or the lepidocrocite (-FeO(OH)) (Figure 12) could produce a similar absorption, and while the 932 nm absorption does not prove the presence of nontronite, it is consistent with the previous CRISM orbital detections by Wray et al. (2009) and Arvidson et al.

(2014). Crucial data for confirming the presence of nontronite in the Matijevic formation is located in the 2300-2400 VSWIR spectral range, which is not included in Opportunity’s PanCam.

Laboratory VSWIR spectra of nontronite- and saponite-bearing samples collected for this study were convolved to MER PanCam VNIR bands and compared to the laboratory identifications (Figure 12). Both of the nontronite-bearing samples have spectral signatures indicative of hematite in the shorter wavelengths (400 – 1100 nm), but no absorption at 932 nm is

visible. In these mixed samples, the Fe-oxides/hydroxides dominate the short wavelengths and mask the nontronite signal. Additionally, two nontronite-lacking samples (CR13 Poás Lake 3 and

CR13 Poás Lake 4 – two montmorillonite- and kaolinite-bearing samples from the bank of Poás’ acidic crater lake) would have been incorrectly identified as containing nontronite, due to the goethite absorptions creating similar spectral minima. Our results indicate that phyllosilicates may be more abundant than can be detected with Opportunity’s PanCam.

2.4.4.3. MSL Curiosity at Gale crater

Mineralogical investigation at Gale crater has relied heavily on Curiosity’s CheMin XRD,

MastCam, and ChemCam’s passive VNIR function (400 – 840 nm), supplemented with elemental chemistry from the Alpha Particle X-Ray Spectrometer (APXS), the Sample Analysis at Mars Gas

Chromatograph (SAM-GC), and ChemCam’s Laser-Induced Breakdown Spectroscopy (LIBS) instrument (eg: Vaniman et al., 2014). This combination of investigative methods provides a wide range of compositional data; however, direct mineralogical data is limited to those materials that are identifiable with XRD, and the narrow spectral window that MastCam and ChemCam’s passive

VNIR sample (400 to 1000 nm). The VNIR broad band sensors on Curiosity face the same challenges and limitations as those utilized on the previously discussed MER Spirit rover. The combination of 400 – 1000 nm VNIR and XRD allows identification of several mineral phases, but underutilizes the VSWIR window’s capabilities and is lacking for phyllosilicate identification.

As shown in this study, XRD is capable of identifying phyllosilicates when they are present in high relative abundance such as the 18 wt % kaolinite in I13 KL cs min2 (Figure 11d) (e.g.,

Vaniman et al., 2014) and down to the 3 wt % kaolinite found in sample Poás Dome Edge 4 (a mixture of primary and secondary minerals collected from the flanks of Poás’ active dome, which was subsequently destroyed in a April 2017 eruption), but fails to identify these materials when

they are present in quantities below ~ 3 wt % (Beckerman, 2016; Swayze et al., 2014). The spectral window covered by MastCam and ChemCam’s passive VNIR is also not sufficient for clear phyllosilicate identification (Table 3, Table 4). It is possible to suggest the presence of these materials using APXS, SAM, and LIBS data, but these analytical techniques provide elemental data, and may not be an accurate reflection of the true mineralogy.

A thorough mineralogical analysis may be conducted using the CheMin XRD, MastCam

VNIR, APXS, and the Sample Analysis at Mars (SAM) Mass Spectrometer (MS) and Gas

Chromatograph (GC), such as that done by Vaniman et al. (2014) on the John Klein and

Cumberland samples at Yellowknife bay. Detailed analysis of XRD data, supplemented with

APXS, VNIR, and SAM-GCMS allowed for the putative identification of a saponitic smectite in the Sheepbed mudstone. However, due to the lack of additional sample processing, the limited detection window (5-50 2) of the CheMin XRD, and the limited spectral window of the

MastCam (400 – 1000 nm), a definitive identification was not possible. Due to these limitations, it is possible that phyllosilicates are far more varied and abundant in Gale Crater than the data suggests. In spite of these instrumental limitations, phyllosilicate characterization at Gale crater remains a key question for the Curiosity rover investigations (Milliken et al., 2016).

Since landing at Gale crater in 2012, Curiosity has analyzed a number of samples that contained or were believed to contain phyllosilicates. The John Klein and Cumberland samples at

Yellowknife Bay (Bristow et al., 2015; Vaniman et al., 2014), the Windjana sample from the

Kimberley Formation (Treiman et al., 2016), the Confidence Hills and Mojave 2 samples in the lower Murray Formation at Pahrump Hills (Cavanagh et al., 2015; Rampe et al., 2016, 2015), and

Telegraph Peak and Buckskin samples farther up the Murray Formation at Pahrump Hills (Morris et al., 2016; Rampe et al., 2016, 2015) were all analyzed via Curiosity’s CheMin XRD. Except

for Telegraph Peak and Buckskin, each of these locations yielded phyllosilicate detections.

However, the relative amounts and type of phyllosilicate vary from one sample to the next

(between 10 – 22 wt % of saponite, collapsed smectite or illite, or kaolinite). The specific types of phyllosilicates remain weakly constrained due to the inability to observe changes in swelling behaviors, or prepare oriented mounts that will increase the distinctive 00l basal reflections, as are commonly utilized in terrestrial laboratory studies. To better assess the ability of Curiosity to identify phyllosilicates in Gale crater, Beckerman (2016) investigated the effectiveness of phyllosilicate identification via rover-analog XRD, and found a discrepancy between the presence of phyllosilicates and the ability to detect them with rover-based XRD methods – specifically noting that the lack of phyllosilicate detections higher up in Pahrump Hills does not prove their absence.

In addition to the challenges of rover-based XRD phyllosilicate identification, there may also be an instrumentation bias for the detection of high temperature SiO2 phases. The discovery of tridymite in the Buckskin mudstone by Curiosity’s CheMin XRD suggests high temperature alteration in or around Gale Crater (Morris et al., 2016). Cristobalite and tridymite – distinct indicators of high temperature systems – must be identified via Curiosity’s XRD due to their relatively featureless spectra across the VSWIR range. This instrumentation bias may cause science investigations to miss other locations containing these deposits – since XRD analyses are limited to locations that the rover can physically access – and could significantly alter our understanding of the geologic history in this region.

2.4.4.4. ExoMars, Mars 2020, and beyond

Future rover investigations would benefit from a combination of hyperspectral VSWIR,

XRD, and Raman. For investigation of phyllosilicate-rich rocks (a common target of interest on

Mars due to the potential for habitability, biosignature preservation, and interpretation of geochemical conditions of formation) VSWIR remains the most effective mineral characterization method. Therefore, the complete VSWIR spectral range (300 – 2500 nm) should be included on future instrumentation to ensure thorough phyllosilicate identification – via the unique metal-OH combinations in the 2100 – 2500 nm range – and to properly utilize reflectance spectroscopy’s strengths. Incorporating a hyperspectral 300 – 2500 nm VSWIR spectrometer on future rover payloads would also allow for identification of the common hydrothermal product anatase, and would add the scientifically valuable ability to ground truth orbital CRISM data. Currently, landing sites for in situ investigations are chosen with an emphasis on orbitally-detected mineralogy. The natural extension of that process should be to match the capabilities of orbital and ground-based VSWIR spectrometers so that we may continue to investigate the presence of geochemically-relevant materials from the surface. By limiting rover-based reflectance spectrometry to broad band imagers, we are significantly crippling our scientific capabilities.

Portions of the VSWIR spectral range will be used for ExoMars’ MicrOmega, Ma_MISS, and

ISEM spectrometers, and the Mars 2020 SuperCam VNIR spectrometer will have partial coverage

(400 – 900 and 1300 – 2600 nm), but the ExoMars PanCam and Mars 2020 MastCam-Z imager will continue to use the 400 – 1000 nm window that has been operated on both the MER and MSL rovers. Most of the spectrometers will operate with improved spectral resolution as well (Table

3). However, the ExoMars PanCam and Mars 2020 MastCam-Z will continue to operate with an

11 band resolution that has been utilized on the MER and MSL rovers. The ExoMars PanCam team has selected filter wavelengths that will allow the instrument to make tentative identifications of sulfates and phyllosilicates (Cousins et al., 2012, 2010) – although these identifications should be confirmed with additional analyses such as XRD. Due to the sensitivity to thin deposits of

Martian dust and the surficial nature of VSWIR sampling, VSWIR observations should be supplemented with additional instruments such as XRD and/or Raman to obtain a more complete mineralogical analysis. It would also be beneficial to all analyses to conduct investigations using surfaces that have been cleared of dust and weathering rinds via onboard brush/drill tools to minimize the effect of surface dust and weathering products.

XRD is best suited for mineralogical analysis of samples rich in sulfates and crystalline

SiO2. However, XRD is lacking when it comes to phyllosilicate identification if present at minor levels, due to rover inability to conduct additional sample processing such as glycolation and the creation of oriented mounts. However, even without these additional processing steps, XRD will continue to be a crucial tool for mineralogical characterization on Mars via the ExoMars Mars-

XRD instrument. XRD is notably absent from the Mars 2020 rover, which may hinder investigations of the landing site and will need to be offset by other instruments included in the payload such as the SHERLOC UV Raman Spectrometer (although this instrument is optimized primarily for the identification of organic materials and not mineralogy), the PIXL XRF spectrometer, and the SuperCam LIBS.

Point analysis via Raman laser spectroscopy results in clean and easy to identify spectra for individual crystals – especially for those samples rich in elemental sulfur, Fe-, and Ti- oxides/hydroxides, which are common on the Mars surface. Although the ExoMars and Mars

2020 Raman spectrometers include high resolution mineral mapping capabilities, this may not be possible for samples with irregular surfaces or large amounts of amorphous materials, phyllosilicates, or Fe-bearing minerals – all of which are likely to be encountered at any future landing site. Point-based Raman analysis with no mineral mapping would very likely miss minor

mineral phases, and should be supplemented with other mineralogical analyses, such as XRD and/or VSWIR.

2.5. Conclusions

We presented the first detailed Mars-relevant instrument comparison used to determine mineralogy from Mars analog hydrothermal deposits. Our results show significant limitations in mineral identification for each instrument: VSWIR, XRD, and Raman. VSWIR, while capable of identifying most hydrothermal materials (with the exception of anhydrous crystalline SiO2 and

TiO2 polymorphs), is extremely sensitive to weathering rinds and other surficial deposits such as

Martian dust. Like the MER rovers, ExoMars and Mars 2020 may circumvent this problem by using onboard drills before conducting contact analyses using any instrument. However, for samples that are not close enough to drill analyses will likely miss a large amount of compositional data that will be obscured by surface dust and weathering rinds.

The other significant limiting factor for Mars-based VNIR is the narrow spectral window that is often utilized due to inherent constraints on instrument capabilities owed to their design

(e.g. Cousins et al., 2012). This narrow window (~400 – 1000 nm for most) is most effective for identification of Fe-bearing minerals. While limited to this narrow window, the ExoMars PanCam band filters have been optimized to provide tentative identifications of other minerals such as sulfates and phyllosilicates, and may also be supplemented by ExoMars’ additional mast-mounted

VNIR spectrometer, ISEM, as well as the MicrOmega and Ma_MISS spectrometers. Although the Mars 2020 Mastcam-Z spectrometer will continue to use the limited 400 to 1000 nm spectral window – which does not cover key regions for the identification of opals (1400, 1900, 2200 nm), sulfates (1400, 1770, 1900, and 2100-2500 nm), zeolites (1400 and 1900 nm), and phyllosilicates

(2100-2500 nm) – the Mars 2020 SuperCam VNIR spectrometer will have detectors to cover both

from 400 – 900 nm as well as 1300 – 2600 nm. This should allow for the detection of the aforementioned minerals – many of which are likely present on Mars and could have important implications for the geologic history and habitability. Expanding the spectral range to 300 – 2500 nm will result in a much more robust and useful tool for mineralogical characterization on Mars compared to previous rover science payloads.

Bulk mineralogical analysis by XRD is exceedingly useful for sample characterization, and is capable of identifying most hydrothermal products that are present above ~3 wt % (Sarrazin et al., 2005). The main shortcoming of XRD is phyllosilicate identification. Although XRD is capable of identifying phyllosilicates, they were often missed in our analysis due to the lack of additional sample preparation. Our preparation-free analysis showed a significant decrease in phyllosilicate detection via XRD compared to VSWIR. This suggests phyllosilicate detection by

MSL Curiosity at Gale Crater may be missing a large amount of clays that could provide new insights into the geochemical and geologic history of the region, and may have significant implications for past habitability.

The addition of Raman Laser Spectroscopy for upcoming Mars investigations does not come without some limitations as well. Raman is not a useful tool for identifying phyllosilicates due to their fluorescent nature, and will therefore not contribute to solving the challenges of a narrow VNIR range and unprocessed XRD samples. Raman is also highly sensitive to elemental sulfur and TiO2 polymorphs – both of which are abundant in hydrothermal systems. These materials easily overwhelm the signal from others, and may result in skewed results. Perhaps the biggest challenge of deploying Raman on Mars is the tendency for Fe-rich materials to burn when in contact with the laser. Fe-rich rocks and Martian dust will provide significant challenges to

Raman data acquisition, but can be overcome.

Future work should be done to address the remaining question of zeolite identification.

The limited number of zeolite-containing samples in this study did not allow for conclusions to be drawn regarding best methods for their identification. These materials are common in basaltic hydrothermal systems, and likely exist on Mars as well (Ehlmann et al., 2009; Wray et al., 2009).

Our results show a Raman laser spectrometer, XRD, and a VSWIR instrument covering the full spectral window from 300 – 2500 nm may be used in tandem to conduct thorough mineralogical assessments of Martian hydrothermal deposits. Each of these instruments has their own weaknesses and challenges – which are often supplemented by the other investigative methods utilized in this study. Current Mars rovers carry some of these instruments, but are likely missing key mineralogical identifications due to the lack of XRD (MER), and limited VNIR spectral ranges (MER and MSL). The upcoming Mars 2020 and ExoMars rovers have robust science payloads planned – including expanded VSWIR spectral ranges. However, the Mars 2020 rover still lacks an XRD instrument, which is vital for bulk mineralogical analyses. Fortunately, the ExoMars rover includes a Raman laser spectrometer, XRD, and multiple VNIR/VSWIR instruments. This combination of instruments should yield detailed mineralogical compositions and provide useful insight into the geochemical conditions on early Mars.

Note: This chapter has been prepared to submit to the Journal of Geophysical Research: Planets as: Black, S. R., Hynek, B. M., McHenry, L. J., McCollom, T. M., Glenister, C., and Cameron, B. I. “Bulk mineralogy of surficial hydrothermal acid-sulfate deposits at Námafjall, Þeistareykir Geothermal Field, and Hengill Volcano, Iceland: Implications for the identification and interpretation of hydrothermal deposits on Mars” Journal of Geophysical Research. All work was completed by SRB. SEM-EDS and XRD alteration product identification was assisted by TMM and LJM. Field ORP values were collected by CG and BIC.

Abstract: The ability of scientists to identify and interpret the geochemical conditions in relict Martian hydrothermal systems is enhanced by study of terrestrial analogs. As previous studies have shown, parent rock composition influences secondary alteration assemblages. Alteration of high Fe Martian basalts must therefore be analyzed through the lens of high Fe terrestrial hydrothermal systems such as those at Námafjall geothermal field, Þeistareykir geothermal field, and Hengill Volcano in Iceland. We characterize the surficial alteration assemblages present at these sites, including three previously undescribed mudpots and a fumarole at Þeistareykir in Iceland’s Northern Volcanic Zone. Results show a wide assortment of SiO2 polymorphs, Fe/Mg/Ca/K/Al-sulfates, Fe/Ti-oxides and hydroxides, phyllosilicates, and zeolites. This includes the identification of 61% Fe natroalunite at Námafjall – adding to a growing list of hydrothermal sites where minerals with intermediate compositions in the alunite-jarosite solid solution have been identified. This suggests intermediate members of the alunite-jarosite series may be common secondary products in hydrothermal systems, and must be further studied in depth to better understand their geochemical limitations and physical characteristics. Our field sites were chosen to span a range of Fe contents in the primary basalts, allowing for investigation of the effects of Fe content in the parent basalts on secondary mineralogy. Our data show a direct correlation between the abundance of primary Fe and secondary Fe-bearing mineralogy. From this trend, we expect alteration of high Fe Martian basalts to produce hydrothermal alteration assemblages with averages of ~10 – 25 wt % Fe-bearing minerals

3.1. Introduction

3.1.1. Hydrothermal systems on Mars

Ever since the Mariner spacecraft returned the first images of Olympus Mons, it has been clear that volcanic processes played a large role in the geologic history of Mars. Volcanoes both large and small are widespread across the surface of the planet (e.g. Crumpler et al., 1996; Wilson and Head, 1994), and several areas have been identified – both from orbital and in-situ data – as relict hydrothermal systems, such as those at Home Plate in Gusev crater (Morris et al., 2008; Ruff,

2015; Ruff et al., 2011; Squyres et al., 2008, 2007; Yen et al., 2008), amorphous silica deposits in

Nili Patera (Skok et al., 2010), mixed sulfates and phyllosilicates in Noctis Labyrinthus (Thollot et al., 2012; Weitz et al., 2011), and Cross crater in Terra Sirenum (Ehlmann et al., 2016). Crater count studies of volcanic features across Mars indicate that intermittent eruptive activity may have spanned much of Mars’ history (Robbins et al., 2011). Impact-induced hydrothermal systems, which have been suggested as an origin for the deposits of chlorite and prehnite in the central peak of the Hesperian-aged Toro crater in Syrtis Major (Marzo et al., 2010), likely occurred across the surface of early Mars (Abramov and Kring, 2005; Rathbun and Squyres, 2002; 2013, 2009). The pervasiveness of volcanic and impact features on Mars combined with widespread evidence of surface and subsurface water (Ehlmann and Edwards, 2014; Hynek et al., 2010; Stuurman et al.,

2016; e.g. James J. Wray et al., 2009) suggests that hydrothermal systems have been common on

Mars, and may have operated throughout much of the planet’s geologic history.

Due to their ability to host thriving microbial communities here on Earth (e.g. Blank et al.,

2002; Walker et al., 2005), ancient Martian hydrothermal systems remain an environment of interest in the search for biosignatures on Mars (e.g. Ruff, 2015). In addition to the ability to host a robust extremophile community that may be similar to primitive terrestrial life forms (Djokic et al., 2017; Reysenbach and Cady, 2001), hydrothermal systems also provide an effective preservation mechanism in the form of opaline silica precipitate (e.g. Al-Hanbali et al., 2001;

Walker et al., 2005) – thereby increasing the probability of biosignature preservation on Mars

(Summons et al., 2011). The ability to host primitive extremophile life forms combined with an effective mode of biosignature preservation makes hydrothermal systems a prime target in the search for signs of life on Mars.

3.1.2 Terrestrial studies of hydrothermal systems

Terrestrial argillic alteration is characterized by strong acid leaching of primary minerals, and the formation of a high silica enrichment zone in the areas proximal to fumarolic vents and hot springs. An alunite-dominated alteration zone often surrounds the silica-enriched core, with a kaolinite-dominated zone in the distal regions (Heald et al., 1987; Pirajno, 2010). Also common are Mg-, Fe-, and Ca-sulfates such as gypsum, jarosite, natrojarosite, Fe-rich natroalunite, hexahydrite, and kieserite, Fe- and Mg- smectites such as nontronite and montmorillonite, elemental sulfur, and Fe-oxides (Bishop et al., 2004; Ehlmann et al., 2012; Hynek et al., 2013;

Marcucci et al., 2013; McCollom et al., 2013a; Rodríguez and van Bergen, 2015).

Several previous investigations have examined the influence of geochemical parameters such as pH, temperature, and fluid:rock ratio on the composition of secondary deposits through the investigation of terrestrial analog sites such as those in Iceland (Table 8). In addition to independent environmental controls such as temperature and fluid:rock ratio, Hynek et al.’s (2014,

2013) survey in Nicaragua, Costa Rica, Iceland, and Hawaii suggest elemental enrichment in the parent lithology provides an abundance of those cations for sulfate formation (eg: alteration of Fe- rich parent rock would result in the formation of Fe-sulfates), and has significant influence over the resulting secondary mineralogies.

Due to its unique combination of a deeply sourced mantle plume (e.g. Wolfe et al., 2002) and mid-ocean ridge volcanism, Icelandic basalts are enriched in Fe – up to ~17 wt % FeOT

(Moune et al., 2007; Nicholson and Latin, 1992; Schuessler et al., 2009; Trønnes, 1990; Trukhin et al., 1995) – and are some of the closest terrestrial matches to Martian basalts, which have been observed to contain up to approximately 21 wt % FeOT in Gusev crater and unaltered Shergottite meteorites (H. Y. McSween et al., 2006; Rubin et al., 2000). Fe enrichment, accessibility, and a

wide variety of volcanic areas and processes distinguish Iceland as a premier location for Mars analog studies.

While many prior studies have sought to characterize various Icelandic hydrothermal deposits (Table 8), our work focuses on the role parent lithology has on the resulting secondary mineral assemblage. We sampled both altered and primary materials at multiple locations across

Iceland (Námafjall, Þeistareykir, and Hengill – geologic setting and previous investigations of which are discussed in sections 3.1.3.1, 3.1.3.2, and 3.1.3.3, respectively), with the goal of collecting altered materials from basalts of varying composition – particularly Fe abundance. In addition to investigating the role of parent lithology and Fe abundance on secondary mineralogy, we also characterize the alteration mineralogy at a new Mars analog site, the Þeistareykir geothermal field in the Northern Volcanic Zone.

Table 8: Previous analog studies of Icelandic hydrothermal acid-sulfate systems Author Location Conditions Alteration products Geptner et al., 2005 Reykjanes, Iceland Low pH fumaroles, Sulfur, hematite, goethite, smectite, pyrite, anatase, hot springs, and lakes, kaolinite/metahalloysite, gypsum, alunogen, bilinite, up to ~80C rhomboclase, jarosite, alunite, boehmite, opal, quartz

Geptner et al., 2007 Námafjall & Þeistareykir Low pH mudpots and Sulfur, anatase, gypsum, Fe-hydroxides and sulfides, Iceland active fumaroles, 10 - 43C zeolites, dioctahedral smectites, kaolinite

Mínguez et al., 2011 Námafjall, Iceland High temperature geothermal Zeolites, smectites (beidellite, nontronite, saponite), field (~ 200C) with low pH kaolinite, halloysite, quartz, calcite, Fe-oxides fumaroles and mudpots

Cousins et al., 2013 Askja & Kverkfjöll, Iceland Glaciovolcanic systems with Sulfur, quartz, palagonite, gypsum, jarosite, pyrite, active low pH fumaroles goethite, hematite, heulandite, smectite 1 - 94C, pH 2 – 7.5

Carson, 2015 Námafjall & Leirhnjúkur, Active low pH fumaroles Amorphous SiO2, cristobalite, quartz, sulfur, gypsum, Krafla region, Iceland and mudpots, halotrichite/pickeringite, epsomite, szomolnokite, ambient – 100C, pH 2 – 2.5 starkeyite, rozenite, rostite, stellerite, alunogen, rhomboclase, (natro)alunite, jarosite, anatase, kaolinite/halloysite, montmorillonite, nontronite, saponite, vermiculite, pyrite/marcasite, hematite, goethite, magnetite, maghemite, clinoptilolite

Harris et al., 2015 Námafjall, Iceland High temperature Opal, sulfur, quartz, calcite, gypsum, alunite, geothermal field with low pH natrojarosite, ferrihydrite, hematite, goethite, anatase, fumaroles and mudpots kaolinite, montmorillonite, nontronite, zeolites

Black and Hynek, 2017 Costa Rica, Iceland, High temperature acid crater Amorphous SiO2, quartz, cristobalite, tridymite, sulfur, Nicaragua lake, active low pH fumaroles, gypsum/anhydrite, (natro)alunite, Fe-rich (natro)alunite, and acid fog regions, (natro)jarosite, rhomboclase, hexahydrite, pickeringite, ambient – 100+C, pH 0 - 4.5 alunogen, hematite, goethite, anatase, kaolinite, montmorillonite, saponite, nontronite, laumontite

3.1.3 Geologic setting of sample sites

Terrestrial analog studies are often used to provide insights into Martian geology and geochemistry. However, the extremely high Fe content observed in Martian basalts (H. Y.

McSween et al., 2006; Ming et al., 2008; Rubin et al., 2000) makes it difficult to find a terrestrial location that is a true geochemical match. The locations chosen for this study – Námafjall and the

Þeistareykir geothermal field in the Northern Volcanic Zone (NVZ), as well as the Hengill

Volcanic Complex in southwest Iceland (Figure 14) – allow us to more closely match the basalt compositions observed on Mars using terrestrial analog sites (Figure 14). Although still not a perfect match for Martian basalts, these locations are more similar to observed compositions of

Mars basalts than the aforementioned commonly-used terrestrial analog sites.

Figure 14: Top: Sampling locations across Iceland; bottom: Fe, Mg, Al, Na, and K oxide abundances measured by XRF and normalized to SiO2 for unaltered Martian basalts in Gusev (H. Y. McSween et al., 2006; Ming et al., 2008) and Meridiani (Rieder et al., 2004), unaltered Shergottite basalts (Hurowitz et al., 2005; Lodders, 1998; Rubin et al., 2000), commonly-utilized terrestrial analog sites in Hawaii (Chemtob and Rossman, 2014; Morris et al., 2000), Nicaragua (Hynek et al., 2013), and Costa Rica (Black and Hynek, 2017), and Icelandic sampling locations for this study.

3.1.3.1 Námafjall geothermal field, Krafla Region

The Námafjall geothermal field is located at 65.64217N, 16.81920W – approximately 8 km south of the main Krafla caldera, 4.4 km east of Lake Mývatn, and along the Krafla rift system running through northern Iceland. Námafjall consists of a ~2-km-long Pleistocene basaltic hyaloclastite ridge (Ármannsson et al., 1987), that has been extensively altered through both fumarolic and mudpot activity. Hyaloclastites form through subglacial volcanic eruptions, and have been proposed to exist on Mars at Olympus Mons (Ann Hodges and Moore, 1979), several low-relief cones near the north polar cap (Garvin et al., 2000), interior layered deposits within

Valles Marineris (Chapman and Tanaka, 2001), and mountainous features within the Dorsa

Argentea Formation near the south polar region (Ghatan and Head, 2002). Alteration of the

Námafjall hyaloclastite ridge provides useful insight into the alteration products that may be present at these and other locations on Mars.

Previous investigation of alteration mineralogy at Námafjall by Harris et al. (2015), Carson

(2015), Geptner et al. (2007, 2005), and Mínguez et al. (2011) used a combination of Visible Near-

Infrared (VNIR) reflectance spectroscopy, X-Ray Diffraction (XRD), Scanning Electron

Microscopy and Electron-Dispersive X-Ray Spectroscopy (SEM-EDS). Identified phases include amorphous and crystalline SiO2, hematite, goethite, kaolinite, smectites, alunite-jarosite group minerals, Ca/Al/Mg-sulfates, anatase, sulfur, and minor amounts of pyrite, ferrihydrite, maghemite, zeolites, calcite and quartz. Alteration at Námafjall is primarily through fluid- dominated mudpots and gas-dominated fumarole vents. At the time of sampling, observed fumarole temperatures were approximately 98C, while mudpot temperatures ranged from 35 to

88C, and pH 1.5 to 2.8. Some mudpots were unsafe to sample directly, therefore pH and temperature ranges may be larger than reported. Sampling locations are shown in Figure 15.

Figure 15: Sampling locations (red stars represent mudpots, and green diamonds indicate fumaroles) at Námafjall. Image courtesy of Google Earth

3.1.3.2 Þeistareykir geothermal field, NVZ

The Þeistareykir geothermal field is located to the north-northwest of Mt. Bæjarfjall, ~23 km north-northeast of Lake Mývatn, and ~4 km west of the Krafla fault system, centered over the

Þeistareykir fissure swarm – the westernmost of the five en echelon fault systems that comprise

Iceland’s NVZ. The most active portion of Þeistareykir geothermal field is on and adjacent to the northern slope of Mt. Bæjarfjall – a subglacial hyaloclastite ridge formed during the last ice age, with surrounding lava flows emplaced approximately 10 – 2.5 kya (Sæmundsson, 2007). While there has been no volcanic activity in the past ~2500 years (Sæmundsson, 2007), the Þeistareykir geothermal field has been identified as one of the highest temperature fields in Iceland, with bedrock temperatures measured up to 380C (Óskarsson et al., 2013), and has persistent hydrothermal surface activity in the form of mudpots and fumaroles.

To date, mineralogical investigations of Þeistareykir have been limited. One study of subsurface alteration in the northeastern portion of the geothermal field (drill site ÞR-7) showed abundant zeolites (laumontite, wairakite, mordenite, and yugawaralite) and smectite to chlorite phyllosilicates (Marosvölgyi, 2009). A surface investigation of the clay-fraction at Þeistareykir by Geptner et al., (2007) identified kaolinite, anatase, gypsum, sulfur, Fe-hydroxides, pyrite, zeolites, and dioctahedral smectites. Other geological investigations at Þeistareykir have focused on aqueous geochemistry and not mineralogy (Ármannsson, 2016, 2004; Ármannsson et al., 1986;

Óskarsson et al., 2013). Variations in pH, temperature, and fluid:rock ratios between adjacent mudpots provide an opportunity to study the effects of these geochemical parameters on the alteration of the same parent material.

Our investigation focuses on surficial alteration by mudpots and a fumarole vent located at

65.87174N, 16.97193W in the Þeistareykir geothermal field (Figure 16). Several mudpots with varying pH (1.5 – 3), temperature (80 – 98C), and varying amounts of fluid are present in the lowlands next to Mt. Bæjarfjall, within the surrounding ~10-2.5 kya lava flows. The nearby fumarole (40C, pH ~0.5) alters a portion of the lowermost flank of the mountain.

Figure 16: Sampling locations at the Þeistareykir Geothermal Field. Image courtesy of Google Earth.

3.1.3.3 Nesjavallavir Power Plant, Hengill Volcanic Complex

The Hengill volcanic region sits atop the rift zone in southwest Iceland, just south of Lake

Þingvellir. Hengill is composed of a central volcano, several smaller shields, and also includes one of the largest geothermal fields in Iceland (Hernández et al., 2012; Marty et al., 1991). Hengill has a complex and active eruptive history, with several postglacial eruptions within the past 10 ky

(Bull et al., 2005; e.g. Foulger and Toomey, 1989; Sinton et al., 2005; Thordarson and Larsen,

2007). Previous investigations of Hengill basalts distinguish them as some of the more Fe-rich in

Iceland (Trønnes, 1990), making them similar to the Gamma Ray Spectrometer (GRS) measurements of the upper 10s of cm of Mars’ crust (Fig. 2) (McSween et al., 2009).

Our investigation at Hengill focused on alteration at two locations on the grounds of the

Nesjavallavir Power Plant (Figure 17). Location 1 (herein referred to as Hengill-1) consists of a fluid-dominated hot spring (pH 2.5 – 3, 77 – 85.9C) located at 64.09722N, 21.27444W. The 90

nearby location 2 (64.08639N, 21.27028W, herein referred to as Hengill-2) consists of a complex fluid-dominated hot spring system with coalescing red (pH 3.5, 55.5C), gray (pH 5.5,

50.0C), and white (pH 6, 22.4C) streams. Dry areas adjacent to the streams at Hengill-2 recorded temperatures up to 88.4C.

Figure 17: Sampling locations (white outlined areas in insets A and B) at the Nesjavallavir Power Plant on the flanks of the Hengill Volcanic Complex. Satellite images courtesy of Google Earth.

Many studies have worked to characterize hydrothermal alteration mineralogy present in the Hengill Volcanic Complex. However, due to Hengill’s large size, these areas are very limited.

Several are in the Hellisheiði Geothermal Field, on the opposite side of Hengill, ~10 km southwest of the Nesjavallavir Power Plant (Eshaghpour, 2003; Gebrehiwot et al., 2010; e.g. Getaneh, 2001;

2015, 2010). In the area around the Nesjavallavir Power Plant, Larsson et al. (2002) and

Kristmannsdóttir and Tómasson (1974) identified oligoclase (secondary feldspar), zeolites, mixed- layer clays (chlorite, Fe-saponite, illite), opal, quartz, prehnite, epidote, pyrite, and hematite in drill cores.

3.2 Methods

3.2.1 Sample collection

Samples from each site were collected during field campaigns in August-September 2016

and June-July 2017. Fresh basalts were collected adjacent to the alteration site to ensure that it

represents the parent substrate. Selection of a wide range of alteration products focused on color

and textural/morphological variations to ensure that all surficial mineralogical products were

collected. Mineralogical sampling at mudpots included both material precipitating directly from

the fluid as well as areas around the perimeter of the mudpot. Mudpot fluid samples were collected

from each site, tested for pH and temperature, and sealed for further analysis.

3.2.2 Analysis

3.2.2.1 Fluid chemistry

Fluids collected from each location were filtered in situ and later acidified with nitric acid and analyzed for trace and rare earth elements and major anions and cations in the Laboratory for

Environmental and Geological Sciences at CU Boulder. Trace elements and REE's were analyzed with a Perkin Elmer SCIEX Elan DRC-e inductively coupled plasma mass spectrometer (ICP-

MS). Indium was used as an internal standard, and four additional standards (blank, 100, 500 and

1000 ppb) were used for calibration. Major cations were measured using an ARL 3410+ inductively coupled optical emission spectrometer (ICP-OES). A blank and three standards were used for ICP-OES calibration. Anions were measured using a Dionex 4500i ion chromatograph using five standards for calibration. An AG14 guard column and AS14 column were used for separation of anions. In situ oxidation-reduction potential (ORP) measurements were made using the Hydrolab multiparameter sonde device from OTT Hydromet. Mineral saturation states (SI)

were calculated using the Geochemists Workbench Spec8 program. Basis entries contained measured ion concentrations, pH, and temperature.

3.2.2.2 Visible Near-Infrared (VNIR) Spectroscopy

Samples were dried at ambient temperature upon return to the lab and VNIR analysis was

completed at Malvern Panalytical using a TerraSpec 4 High Resolution Reflectance Spectrometer.

The TerraSpec 4 spectrometer measures reflected light between 350 and 2500 nm with a 6 nm/band

spectral resolution. A halogen light source is located inside the ~ 2 cm diameter contact probe to

eliminate effects from outside light sources. The spectrometer was white referenced and calibrated

regularly throughout data collection to ensure high data quality and signal to noise ratio.

Reflectance spectra were analyzed using the ENvironment for Visualizing Images (ENVI)

software. Laboratory spectra of field samples were matched to several reference libraries,

including the USGS splib06 (Clark et al., 2013), the CRISM spectral library (PDS Geosciences

Node and CRISM Spectral Library Working Group, 2014), and McCollom et al.’s (2014)

natroalunite-natrojarosite solid solution series. Continuum-removed spectra were used for more

accurate band depth and center location.

VNIR reflectance spectroscopy may be used to identify a wide range of materials, and is

particularly useful for characterizing phyllosilicates and surficial deposits such as weathering rinds

(Black and Hynek, 2017; Marcucci et al., 2013). It is important to note, however, that VNIR is

exceptionally sensitive to surface coatings and does not “see” underlying materials (Black and

Hynek, 2017; Marcucci et al., 2013). Therefore, it is necessary to use VNIR in conjunction with

additional investigative methods such as bulk powder XRD and Raman laser spectroscopy to

thoroughly characterize sample mineralogy.

3.2.2.3 X-Ray Diffraction (XRD)

After drying, all samples were ground to ~ 10 m using an agate mortar and pestle, and then analyzed with the Terra field-portable XRD from Olympus Scientific Solutions (Olympus

Scientific Solutions, 2014). The Terra XRD utilizes a Cu-K x-ray source ( = 1.5418 Å), with detectors from 5 - 55 2, a ~0.05 2 step size, and 45 minute integration time. The Terra XRD can identify crystalline materials that are present at ~1 wt % or higher (Treiman et al., 2010).

Resulting diffraction patterns were processed using the XPowder software, the American

Mineralogist Crystal Structure Database “difdata” library, McCollom et al.’s (2014) natroalunite- natrojarosite solid solution series diffraction patterns, and Jacobsen et al.’s (2014) Oskarssonite

(AlF3) diffraction patterns. Quantitative mineral abundances were calculated using the Reference

Intensity Ratio (RIR) tool in XPowder after matching the sample to library diffraction patterns.

3.2.2.4 X-ray Fluorescence (XRF)

Fresh basalt samples were analyzed by X-ray Fluorescence (XRF) to determine their major and minor element compositions. Samples were powdered using a shatterbox, with 0.5 g of powdered sample roasted in a high temperature oven at 1,050 °C to determine the loss-on- ignition

(LOI) of the sample. 1.0000 g (+/- 0.0003 g) of powdered material was mixed with 10.0000 g (+/-

0.0003 g) of lithium tetraborate flux to create a 1:10 ratio to fuse into glass beads using a Claisse

M4 fluxer. The resulting glass bead was analyzed for major and minor elements using a Bruker S4

Pioneer WD-XRF following the methods of McHenry (2009).

3.2.2.5 Electron Microprobe Wavelength-Dispersive Spectrometer (EPMA-WDS)

Polished thin sections of fresh basalts were carbon-coated and analyzed for primary mineral composition using a JEOL JXA-8230 Superprobe wavelength-dispersive spectrometer

(WDS). Major element compositions (Si, Fe, Mg, Ca, Na, K, Ti, F, O) of 5-10 examples of each primary mineral type (feldspar, pyroxene, olivine, and glass – when present) were measured for

each fresh basalt sample. Measured compositions were then averaged to obtain an average mineral

composition for each parent basalt.

3.3 Results

3.3.1 Primary basalt composition

3.3.1.1 Námafjall geothermal field, Krafla Volcano

Due to the porous and friable nature of hyaloclastites, unaltered basalts are difficult to find

at Námafjall. Petrographic and EPMA-WDS analysis of polished thin sections reveal abundant

phenocrysts of glass, plagioclase feldspar (An76), and clinopyroxene (diopside), as well as small

vesicles throughout the sample. No original olivine is present, and the originally glassy matrix is

extensively altered to a 14.7 Å smectite clay even in the freshest of samples, as seen via X-Ray

Diffraction (XRD). X-Ray Fluorescence (XRF) analysis place Námafjall at the high end of the

SiO2 range, with moderate FeOT and low MgO relative to Þeistareykir and Hengill (Table 9). It is

important to note that due to the loss of original olivine, and highly altered glassy matrix at

Námafjall, these bulk elemental abundances may not accurately reflect the original composition

of the parent basalt. However, compositions of fresh Námafjall hyaloclastites were also reported

by Geptner et al. (2007) and are similar to our collected samples. This suggests that despite the

present alteration, these may be considered representative compositions of parent basalts at

Námafjall.

Table 9: Elemental weight percent oxide compositions measured with XRF for fresh basalts collected in this study SiO2 Al2O3 FeOT MgO MnO CaO Na2O K2O TiO2 P2O5 Total Námafjall 48.5 14.6 10.1 5.9 0.2 11.8 1.9 0.3 1.9 0.2 95.4 Þeistareykir: fumarole 47.4 13.9 11.5 8.6 0.2 11.5 1.8 0.2 1.6 0.2 96.9 Þeistareykir: mudpots 48.1 14.7 9.3 11.5 0.2 12.2 1.3 0.1 0.7 0.1 98.2 Hengill 46.5 14.4 11.8 6.6 0.2 12.2 1.9 0.2 2.2 0.2 96.2

3.3.1.2 Þeistareykir geothermal field, NVZ

Samples of fresh basalts collected from lava flows adjacent to Mt. Bæjarfjall (Þeistareykir: mudpots) consist of large vesicles and olivine phenocrysts (Fo89), clinopyroxenes (diopside), plagioclase feldspar (An79), and a groundmass consisting primarily of plagioclase and clinopyroxene microlites. XRF analysis of these parent basalts shows high SiO2, MgO, and Al2O3, with lower FeOT relative to the adjacent Mt. Bæjarfjall basalts (Þeistareykir: fumarole), as well as basalts from Námafjall and Hengill (Table 9).

Fresh basalts gathered at the Þeistareykir fumarole location contain large phenocrysts of plagioclase feldspar (An76), clinopyroxene (hedenbergite), and olivine (Fo81), as observed in thin section and EPMA-WDS analysis. The groundmass consists of glass, plagioclase feldspar, and clinopyroxene microlites. Although located on the flanks of the Mt. Bæjarfjall hyaloclastite, the surface basalt being altered at this fumarole is competent, does not contain vesicles, and may be sourced from capping lava flows that have eroded from the uppermost portion of Mt. Bæjarfjall.

These Þeistareykir basalts contain moderate SiO2 and FeOT, lower Al2O3, and higher MgO relative to Námafjall and Hengill (Table 9).

3.3.1.3 Nesjavallavir power plant, Hengill Volcanic Complex

Previous investigations of Hengill basalts yielded FeOT abundances of up to ~17 wt %

(McHenry et al., 2016; Trønnes, 1990). Basalts collected at Hengill-1 and Hengill-2 appear to have lower FeOT; although this is still higher than many terrestrial basalts. Basalt substrates for these sites also have the lowest SiO2 of those investigated for this study (Table 9). Petrographic and EPMA-WDS analysis of polished thin sections show large phenocrysts of plagioclase feldspar

(An80), clinopyroxene (diopside), olivine (Fo79), and possibly minor orthopyroxene (tentatively identified in petrographic thin section, but not confirmed by other methods). The matrix is

extremely fine, and composed of glass, plagioclase feldspar, and clinopyroxene microlites. No vesicles are present.

3.3.2 Secondary mineralogy

Secondary mineralogy at each site was classified by major mineral components. Most samples were dominated by one or two minerals that were present at > 20 wt % in the bulk XRD pattern. Those which were present at 10 – 20 wt % were identified as major components, and 5 –

10 wt % identified as minor components. Minerals present at < 5 wt % were deemed trace components. Minerals that were identified by VNIR but not present in bulk XRD were assumed to be present below the XRD detection limit of ~1 wt %, and were also included in the trace component category.

3.3.2.1 Námafjall geothermal field, Krafla Volcano

VNIR spectra and XRD patterns of alteration products at Námafjall’s mudpots reveal amorphous SiO2, hematite, anhydrite, gypsum, pickeringite, rhomboclase, voltaite, unidentified

Al-sulfate, Al-phyllosilicates, and elemental sulfur (Table 10, Figure 18). Materials sampled include precipitates as well as mud/sediment around the edges of the mudpots. The average bulk mineralogy at the two mudpot areas sampled at Námafjall (indicated in Figure 15) as determined by XRD shows abundant Ca- and Fe-sulfates, elemental sulfur, amorphous SiO2, hematite, and

Al-phyllosilicates. Bulk XRD of samples collected from fumarole locations at Námafjall also reveals ~60% Fe natrojarosite (Fe % indicates the relative amount of Fe present in the B site in the alunite group formula AB3(SO4)2(OH)6), (natro)alunite, apjohnite/halotrichite (as the two are nearly indistinguishable in XRD, here we will classify them as halotrichite group sulfates), Al- phyllosilicates, and smectite. VNIR identified additional amorphous SiO2, hematite, nontronite, and an unspecified Mg-sulfate in the fumarole samples, and additional hematite, montmorillonite,

and amorphous SiO2 in the mudpot samples. The presence of these materials in VNIR reflectance spectra but not XRD patterns suggests they are present in thin weathering rinds, and have relative bulk abundances below the XRD detection limit of ~1 wt %.

Hydrothermal fluid gathered from a hot spring on the flanks of Námafjall contains higher amounts of most elements relative to our other study locations. Námafjall fluids are enriched in

Si and Fe, with Si > Fe > Al > Ca > Mg > Na > K > Mn > Ti, and is depleted in Cl- relative to

Hengill and Þeistareykir (Table 11). Saturation states (log (Q/K), where Q = the reaction quotient and K = the equilibrium constant) were calculated for common hydrothermal products, with results

(supersaturated, saturated, or undersaturated) shown in Table 11. Minerals with a log (Q/K) = 0 are considered saturated. For our calculations, we consider a log (Q/K) from -0.4 – 0.4 to be saturated to account for error in thermodynamic databases.

Table 10: Secondary alteration mineralogy identified with VNIR and XRD at Námafjall geothermal field Dominant Major Minor Trace Sample Setting (> 20 wt %) (10 - 20 wt %) (5 - 10 wt %) (< 5 wt %) Nama U1 Gy An Hem* Nama U2 high F:R S Mo *, Am* Nama U3 mudpots; Gy Am Hem* Nama U10 35 – 90.6C Al-phy Hem BMH N1 biomin pH 1.5 - 3 Am, Gy, S BMH N2 Al-sulf Rho, Vol Pk, S Nama U4 Fe61-Alu, Na-Alu, Am Gy Nama U5 Kaol Alu Nama U6 Gy An Hem* Nama U7 Gy An Hem* Nama U8 GRAD 1 S, Alu Am* Nama U8 GRAD 2 Al-phy Alu Mg-sulf* Nama U8 GRAD 3 Al-phy Mg-sulf*, Hem* Nama U9 orange Fumarole Am Hem, Ep, Jar, Rho Nama U9 red ~ 98C Am, Hem Ana, Rho Nama U9 white Am, Ana, Rho BMH N3 grad 1 S Am* BMH N3 grad 2 Am, S, Ana BMH N3 grad 3 Am, Ana, Min Hem BMH N4 S, Ana BMH N5 Ap/Hal Roz BMH N6 Sm (VNIR: Non, Mo)

Al-phy – Al-phyllosilicate; Al-sulf – unidentified Al-sulfate; Am – amorphous SiO2; Alu – alunite; Ana – anatase; An – anhydrite; Ap/Hal – apjohnite/halotrichite; Ep – epsomite; Fe61-Alu – 61% Fe natrojarosite; Gy – gypsum; Hem – hematite; Jar – jarosite; Kaol – kaolinite; Mg-sulf – Mg-sulfate; Min – minamiite; Mo – montmorillonite; Na-Alu – natroalunite; Non – nontronite; Pk – pickeringite; Rho – rhomboclase; Roz – rozenite; Sm – smectite; S – sulfur; Vol – voltaite; * indicates minerals which were identified with VNIR but were not present in XRD and are therefore assumed to be present below the XRD detection limit of ~1wt %.

Figure 18: Representative samples from Námafjall with annotated XRD patterns and VNIR reflectance spectra; Far right: cm-scale dried hand samples; Al-phy – Al-phyllosilicate; Al-sulf – Al-sulfate; SiO2 – amorphous SiO2; An – anhydrite; F – primary feldspar; Fe – iron; 61 – ~60% Fe natrojarosite; Gy – gypsum; Hal – halotrichite group; H – hematite; K – kaolinite; Mo – montmorillonite; Na – natroalunite; Non – nontronite; P – primary pyroxene; Pk – pickeringite; R – rhomboclase; Roz – rozenite; Smec – smectite; S – sulfur; V – voltaite

Table 11: Fluid geochemistry: measured ion concentrations and saturation states for common hydrothermal minerals, and field ORP values Hengill-1 Hengill-2 Námafjall Þeistareykir pH 2.5 3.5 2 3 Measured ion concentrations (M) NO2 n.d. n.d. n.d. n.d. Br 0.00000221 0.00000340 n.d. n.d. F 0.0000115 0.00000993 0.0000138 0.00000428 NO3 0.0000714 0.00000755 0.0000218 0.0000126 K 0.0000675 0.0000144 0.0000409 0.0000101 Cl 0.000371 0.000148 0.00000818 0.0000152 Na 0.000671 0.000389 0.000216 0.0000664 Ca 0.000893 0.000736 0.000913 0.000242 Mg 0.00132 0.000533 0.000857 0.000393 PO4 0.0000159 0.00000715 0.00000629 0.00000674 Mn 0.0000247 0.00000899 0.0000160 0.00000441 SO4 0.0102 0.00230 0.0129 0.0159 Ti 0.000000299 0.000000131 0.000000588 0.000000527 Fe 0.00280 0.0000406 0.00176 0.000471 Al 0.00113 0.000143 0.00282 0.00185 Si 0.357 0.201 0.495 0.458 Mineral saturation states Nontronite, clinoptilolite, Nontronite, clinoptilolite, Nontronite, clinoptilolite, Nontronite, clinoptilolite, pyrophyllite, hematite, pyrophyllite, hematite, pyrophyllite, hematite, pyrophyllite, hematite, Supersaturated heulandite, mordenite, heulandite, mordenite, heulandite, mordenite, heulandite, mordenite, log (Q/K) > 0.4 quartz, tridymite, goethite, quartz, tridymite, goethite, quartz, tridymite, goethite, quartz, tridymite, goethite, cristobalite, kaolinite, cristobalite, kaolinite, cristobalite, kaolinite, cristobalite, SiO2(am) SiO2(am), jarosite, illite SiO2(am) SiO2(am), jarosite Saturated Alunite - Illite Kaolinite -0.4 < log (Q/K) < 0.4 Observed ORP values* 291 96 – 359 93 - 477 203 – 495 n.d = not detected * Low ORP values (< 200) indicate reducing conditions; > 250 indicates strongly oxidizing conditions

3.3.2.2 Þeistareykir geothermal field, NVZ

The high fluid:rock mudpots 1 and 2 displayed similar mineral assemblages (Table 12,

Figure 19), dominated by elemental sulfur, gypsum, goethite, halotrichite group minerals, pyrite, and Al-phyllosilicates such as kaolinite. Additional components include jarosite, hematite, and minor amounts of amorphous SiO2. Additional goethite and amorphous SiO2 were detected using

VNIR and have estimated relative abundances in those bulk samples of  1 wt %. Samples collected at the adjacent low fluid:rock mudpot 3 contain primarily amorphous SiO2, elemental sulfur, Fe- and Ti-oxides/hydroxides, and kaolinite. Sulfate abundance is significantly lower than at mudpots 1 and 2, and contain roughly equal percentages of Fe-, Al-, and Mg-sulfates. No Ca-

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sulfates were detected at mudpot 3, and additional kaolinite and montmorillonite were identified in VNIR spectra. Fumarole mineralogy (sampling location indicated in Figure 16) at Þeistareykir is dominated by amorphous SiO2, Ca-, Al-, and Fe-sulfates, Al-phyllosilicates, and vermiculite.

Cristobalite, elemental sulfur, and Fe- and Ti-oxides/hydroxides are also present. Additional montmorillonite and amorphous SiO2 were identified in VNIR spectra. Fluids collected at mudpot

#1 (high fluid:rock) show moderate levels of most major ions relative to Hengill and Námafjall, with Si > Al > Fe > Ca > Mg > Na > K > Mn > Ti, with an absence of Br- (Table 11).

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Table 12: Secondary alteration mineralogy identified with VNIR and XRD at Þeistareykir geothermal field Dominant Major Minor Trace Sample Setting (> 20 wt %) (10 - 20 wt %) (5 - 10 wt %) (< 5 wt %) P1 Kaol Hem, Goe*, Am* P2 Al-phy Goe*, Am* P3 S Kaol Hem, Rho High F:R P5 Al-phy Am* mudpots P7 Mo Goe, Jar Am* (MP 1 & 2) P8 Gy, Mo Goe* ~ 80C P9 Gy, Mo Goe*, Am* pH 3 BMH P1 S Kaol BMH P2 Al-phy S, Py P10 Ap/Hal, Kaol P11 S Rho Am* P12 Am, Hem S, Rho Mo* P13 Kaol Goe*, Am* P14 Am, S Kaol Mo* P15 Am, Goe Hem, Ana Mo* P16 red Hem P16 yellow Am, Kaol Hem P17 purple Hem Kaol*, Am* P17 yellow-green Hem Jar Kaol*, Am* P18 blue Low F:R Kaol, Hem Am* P18 green mudpot Al-phy Hem P18 red (MP 3) Am, Kaol, Hem P18 white ~ 98C Kaol P19 white pH 1.5 S Am* P19 yellow S Am* P20 blue Am, Hem Mo* P20 orange Goe Am* P20 red Am Hem, Ep, Mir P20 tan Am Hem* P20 white Am, Kaol P20 purple Am, Hem BMH P3 Hem Mo* BMH P4 Am, Hem Kaol* Pfum 1 Ap/Hal Cri, Rho Pfum 2 Na-Jar, Ver, Gy, Ana Mo*, Am* Fumarole Pfum 3 Am, S Gy, Al-phy, Alu ~ 40C Pfum 4 S pH ~ 0.5 Pfum 5 Gy, Ver S Mo* BMH P6 min Gy Pk, Alg Am

Al-phy – Al-phyllosilicate; Alu – alunite; Alg – alunogen; Am – amorphous SiO2; Ana – anatase; Ap/Hal – apjohnite/halotrichite; Cri – cristobalite; Ep – epsomite; Goe – goethite; Gy – gypsum; Hem – hematite; Jar – jarosite; Kaol – kaolinite; Mir – mirabilite; Mo – montmorillonite; Na-Jar – natrojarosite; Pk – pickeringite; Py – pyrite; Rho – rhomboclase; S – sulfur; Ver – vermiculite; * indicates minerals which were identified with VNIR but were not present in XRD and are therefore assumed to be present below the XRD detection limit of ~1wt %.

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Figure 19: Representative samples from Þeistareykir with annotated XRD patterns and VNIR reflectance spectra; Far right: cm-scale dried hand samples; Al-phy – Al-phyllosilicate; Alg – alunogen; SiO2 – amorphous SiO2; At – anatase; C – cristobalite; Ep – epsomite; Fe – iron; G – goethite; Gy – gypsum; Hal – halotrichite group; H – hematite; J – jarosite; K – kaolinite; Mo – montmorillonite; Pk – pickeringite; Py – pyrite; R – rhomboclase; S – sulfur; Vr – vermiculite 103

3.3.2.3 Nesjavallavir power plant, Hengill Volcanic Complex

Both locations sampled at Hengill are fluid-dominated hot spring systems. Although their secondary mineral assemblages are similar (Table 13, Figure 20), the relative abundance of mineral phases vary. Both locations have abundant amorphous SiO2, sulfur, sulfates, Fe- and Ti- oxides/hydroxides, and Al-phyllosilicates. Oxides/hydroxides at Hengill-1 are primarily anatase, with minor ( 1 wt %) goethite identified in VNIR. Sulfates at Hengill-1 are predominantly the halotrichite group, hexahydrite, gypsum, alunite, and minor amounts of jarosite. Montmorillonite is also abundant at Hengill-1. Hengill-2 alteration products also contain large amounts of amorphous SiO2, elemental sulfur, sulfates, oxides/hydroxides, and Al-phyllosilicates, however, oxides/hydroxides are more abundant overall relative to Hengill-1, and are nearly equal parts Fe- bearing (hematite and goethite) and anatase (TiO2). Sulfates at Hengill-2 consist primarily of alunogen and alunite (Al-sulfates), with moderate amounts of jarosite. No Ca-sulfates were identified at Hengill-2. Several Al-phyllosilicate samples were identified with XRD, with two other kaolinite-bearing samples identified in VNIR spectra, along with 3 montmorillonite-bearing samples.

Fluids collected from the Hengill-1 and Hengill-2 locations have markedly different compositions (Table 11). Both fluid samples contain similar levels of Br-, however, Hengill-2 fluid (collected from the red stream) is significantly depleted in all other major ions relative to

Hengill-1. Stream beds at the Hengill-2 site contain thick biofilms which likely impact the fluid

2- 3+ 2- 4+ 3+ 3+ 4+ chemistry. Both Hengill fluid samples are depleted in PO4 , Mn , SO4 , Ti , Fe , Al , and Si relative to Námafjall and Þeistareykir.

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Table 13: Secondary alteration mineralogy identified with VNIR and XRD at Nesjavallavir power plant, Hengill Dominant Major Minor Trace Sample Setting (> 20 wt %) (10 - 20 wt %) (5 - 10 wt %) (< 5 wt %) H1 crenulated white Qtz, Gy S, Ac Am* H1 gray Al-phy Hx, S Mo*, Am* H1 orange High F:R Mo Jar Goe*, Am* H1 peach mudpot Am, Ana S H1 red 77 - 86C Al - phy Alu Mo* H1 white pH 2.5 - 3 S Mo* H1 yellow Am, Kaol BMH H1 Ap/Hal H2 crenulated Alg, Kaol, Qtz H2 gray S Mo* H2 gray 2 Am Ana, S, Alu Jar, Py H2 orange Kaol H2 peach rind Am, Ana Mo* H2 purple High F:R Am, Hem, Goe Kaol*, Mo* H2 red hot spring Goe Kaol* H2 white 22 - 55C Am Ana, S, Py Kaol H2 white rind pH 3.5 - 6 Al-phy Ana Am* H2 white 2 S H2 yellow S, Ana Am* BMH H2 Al-phy Goe Am* BMH H3 Al-phy, S, Alu Hem, Py BMH H4 Al-phy Hem Mo*, Goe*

Al-phy – Al-phyllosilicate; Alu – alunite; Alg – alunogen; Am – amorphous SiO2; Ac – analcime; Ana – anatase; Ap/Hal – apjohnite/halotrichite; Goe – goethite; Gy – gypsum; Hem – hematite; Hx – hexahydrite; Jar – jarosite; Kaol – kaolinite; Mar – marcasite; Mo – montmorillonite; Py – pyrite; Qtz – quartz; S – sulfur; * indicates minerals which were identified with VNIR but were not present in XRD and are therefore assumed to be present below the XRD detection limit of ~1wt %.

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3.4 Discussion

3.4.1 Secondary mineralogy

3.4.1.1 Námafjall geothermal field, Krafla Volcano

Our investigation of alteration assemblages at Námafjall has identified several new minerals that were not present in previous studies (Carson, 2015; Geptner et al., 2007; Harris et al., 2015; Mínguez et al., 2011). ~60 % Fe natroalunite, minamiite, voltaite, and halotrichite group minerals were all present in our Námafjall samples, adding to the diverse mineralogy identified at this site. The presence of Al-phyllosilicates at Námafjall’s mudpots indicates a high fluid:rock ratio, as well as an average pH of ~ 3 to 4 (Hynek et al., 2013; Jercinovic et al., 1990; Paul and

Zaman, 1978). Observed pH was below this range, representing what is likely a brief fluctuation to more acidic conditions at this site, or the Al-phyllosilicates being stable locally in more acidic conditions. This average pH ~3 – 4 would result in leaching of Ca, Na, K, and Mg from the surrounding parent rocks, while Al, Si, and Fe remained as immobile components of the system

(Banin et al., 1997; Hurowitz et al., 2006; Jercinovic et al., 1990; Paul and Zaman, 1978; Thorseth et al., 1991) forming the observed Al-phyllosilicates, Fe-sulfates and oxides/hydroxides, and amorphous SiO2 (Table 11). Al-phyllosilicate-rich locations are indicative of advanced argillic alteration and extreme leaching (Pirajno, 2010), and may indicate areas where hydrothermal fluid was channeled, and the strongest alteration was occurring. One sample (I16 Nama U2) contained trace amounts of montmorillonite, and may indicate an area of localized increased pH (Berger and

Velde, 1992), possibly due to a lower fluid:rock ratio or fluid dilution by precipitation.

Alternatively, the montmorillonite found at this location may be forming at a pH outside the range that is currently accepted for montmorillonite formation. At the time of sampling, fluid temperatures ranged from 35 – 90.6C. The presence of Al-phyllosilicates indicates a time of

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slightly elevated temperatures (100 – 120C) (Pirajno, 2010). Abundant gypsum deposits likely formed from evaporation of SO4-rich hydrothermal fluids as they reached the surface, enriched in the highly mobile Ca, which was leached from parent rocks below. Many gypsum deposits are also visible as large diagenetic veins throughout the site. While the gypsum collected in our samples was present in large apron deposits and is not from the veins themselves, these veins are likely a source of additional Ca that may be incorporated into the hydrothermal fluid and contribute to the large gypsum aprons around mudpots and fumaroles at this location.

Samples collected from dry areas around fumaroles contained large amounts of elemental sulfur, amorphous SiO2, anatase, natroalunite, ~60% Fe natroalunite, rozenite, and Al- phyllosilicates, which reflects deposition of S0 from a vapor phase and residual enrichment of Si,

Al, Fe, and Ti by acid leaching at low pH (Banin et al., 1997; Hurowitz et al., 2006; Jercinovic et al., 1990; Paul and Zaman, 1978; Thorseth et al., 1991). Abundant nontronite and montmorillonite indicate temperatures up to ~ 200C in the fumarole areas (Barth-Wirsching and Holler, 1989;

Henley and Ellis, 1983; Pirajno, 2010). Gypsum/anhydrite, halotrichite group minerals, and trace amounts of Mg-sulfate likely formed through evaporative processes, forming surficial deposits.

The presence of ~60% Fe natroalunite is the fourth detection of a naturally-occurring intermediate member of the alunite-jarosite solid solution series. The presence of intermediate Fe-rich natroalunite was identified by SEM-EDS analysis from McCollom et al., (2013a). Intermediate alunite-jarosite members were also reported by McCollom et al. (2014) at fumarolic sites in

Nicargua and by Beckerman (2016) and Black and Hynek (2017) at Poás and Turrialba volcanoes in Costa Rica. The presence of naturally-occurring intermediate members of the alunite-jarosite group has implications for the detection and interpretation of Martian deposits (McCollom et al.,

2013, 2014), and is further discussed in section 3.4.2.2.

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In addition to individual samples, three-sample gradients were taken from two separate fumarole vents. The BMH N3 gradient (Figure 21a) transitions from nearly pure elemental sulfur nearest the fumarole vent (0.40 m, 98C) to amorphous SiO2, anatase, minamiite, and hematite

1.22 m from the vent (60C) (El-Maarry et al., 2017). The Nama U8 gradient samples (Figure

21b) show a transition from elemental sulfur and alunite proximal to the vent (0.30 m, 97.2C) to primarily Al-phyllosilicate 1.22 m from the vent (31.6C), with trace amounts of an unidentified

Mg-sulfate and hematite (El-Maarry et al., 2017). With the exception of vent-proximal alunite formation in our Nama U8 gradient, which may also coexist with SiO2 in advanced argillic alteration zones (Pirajno, 2010), gradient mineralogy is consistent with those found by Hynek et al. (2013) in Nicaraguan volcanoes and is a reflection of variations in temperature and fluid composition. In the areas immediately surrounding the fumarole vents, all cations have been leached except for Si and Al, creating a SiO2 and alunite rich zone, with elemental sulfur precipitating directly from the vapor phase. The degree of cation leaching decreases with distance from the vent, leaving Fe, Mg, and Ti in addition to Al and Si, resulting in more diverse mineralogical mixtures.

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Figure 21: Mineralogy of Námafjall gradient samples BMH N3 (left) and Nama U8 (right); stars indicate vent location

3.4.1.2 Þeistareykir geothermal field, NVZ

Previous investigations of alteration mineralogy at Þeistareykir have been limited to a small number of studies, with focused on the clay fraction (Geptner et al., 2007), or were characterizing subsurface alteration in a different portion of the Þeistareykir geothermal field (Marosvölgyi,

2009). Our investigation of surface alteration mineralogy at Þeistareykir is the first to characterize both the clay faction and coarse crystalline components within our study area. Although the

Þeistareykir mudpot sampling locations are adjacent and altering the same parent material, the different fluid:rock ratios and resultant fluid chemistry result in dramatically different alteration assemblages (Table 12, Figure 22). Despite present-day surface ORP values of 203 – 495 mV

(moderately to strongly oxidizing) for the high fluid:rock mudpots 1 and 2, the abundance of pyrite in these mudpots suggest more reducing conditions in the subsurface. Alteration mineralogy at

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the low fluid:rock mudpot 3 (pH 1.5, 60 – 98C) contained large deposits of surficial hematite, with a moderately oxidizing ORP of 231 mV.

Significant amounts of Fe-, and Al-bearing minerals (pyrite, goethite, hematite, halotrichite group minerals, and Al-phyllosilicates) suggest enrichment of Fe and Al either by leaching of other ions (Ca, Na, K, Mg), or for Al, pH fluctuations resulting in mobilization (pH < 3) and re- precipitation (pH > 5) of Al-bearing minerals (Banin et al., 1997; Drever, 1997; Hurowitz et al.,

2006; Jercinovic et al., 1990; Paul and Zaman, 1978; Stumm and Morgan, 1996; Thorseth et al.,

1991). The presence of jarosite indicates the pH at this location was < 4 (Klingelhöfer et al., 2004;

Stoffregen, 1993), suggesting leaching of other elements at pH 3 - 4, with Fe and Al remaining immobile. There is also the potential of more Fe precipitating as deep-sourced low pH hydrothermal fluids reach the surface and the continued addition of leached ions to the solution

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buffers the pH to just above pH 3. After being leached from the surface parent basalt, large amounts

of Ca must have re-entered the system to form the gypsum deposits seen here. This was likely

through surface evaporation of hydrothermal fluids which were enriched in Ca, leached from

parent rocks below. The presence of pyrite indicates reducing conditions just below the surface,

while goethite and hematite are present in the oxidized surface layer. Interspersed kaolinite and

montmorillonite may indicate areas of varying pH - kaolinite present in areas where fluids are

channeled and lower pH is maintained by constant replenishment of fluids, while montmorillonite

locations may show areas with lower fluid replenishment rates, allowing for moderate buffering

and a slight increase in pH.

Although mudpots 1 & 2 and mudpot 3 have many secondary mineral phases in common

(Table 12), there are several distinct differences. Amorphous SiO2 is noticeably absent from the high fluid:rock mudpots, only identified in trace amounts with VNIR. The lack of SiO2 precipitation implies very limited Si availability for this system, however, analysis of fluids collected from mudpot 1 show 0.458 M Si (Table 11), which is well above the levels at which SiO2 should precipitate (Iler, 1979; Jonckbloedt, 1998). This is suggestive of kinetic inhibitors for SiO2 precipitation in the high fluid:rock mudpots, and requires further investigation of this site for a more detailed analysis. While mudpot 3 is currently a low fluid:rock environment with observed ground temperatures from 60 – 98C, the presence of coarse-grained gray (specular) hematite found on several boulders scattered across the surface (Figure 23), and abundant phyllosilicates indicates higher fluid:rock and temperature (> 100C) conditions in the past (Pirajno, 2010). Although gray hematite often forms by oxidation of magnetite in hydrothermal fluids, it may also form by precipitation from a vapor phase (eg: 2FeCl3 (g) + 3H2O (g) → Fe2O3 (s) + 6 HCl (g)) (Catling and

Moore, 2003). However, no HCl was detected in fluid samples from this location, rendering this

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formation pathway unlikely. In addition to gray hematite, several samples contained fine grained red hematite, which is consistent with the present day low fluid:rock ratio and strongly oxidizing conditions (Langmuir, 1971; Torrent et al., 1982; Tosca et al., 2008). There is also a complete lack of gypsum at mudpot 3, suggesting either no secondary Ca input from hydrothermal fluids to form evaporite gypsum crusts as are seen at mudpots 1& 2, or if the gypsum crusts formed, they have since been removed – either through weathering and erosion, or dissolution at a time of higher fluid:rock ratios.

Figure 23: Coarse-grained gray (specular) hematite at Þeistareykir mudpot 3

Samples collected at the Þeistareykir fumarole were dominated by elemental sulfur, amorphous SiO2, gypsum, halotrichite group minerals, Al-phyllosilicate, and vermiculite (Table

12). Gypsum deposits at this location likely formed from the evaporation of small amounts of acidic steam after leaching Ca from the parent rock. Despite the present fluctuation of pH to ~0.5, sustained leaching at pH 3 – 4 would result in enrichment of the immobile Si, Ti, Fe, and Al in the surrounding rock (Banin et al., 1997; Hurowitz et al., 2006; Jercinovic et al., 1990; Paul and Zaman,

1978; Thorseth et al., 1991), and formation of the observed amorphous SiO2, cristobalite, Al-

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sulfates (alunite and halotrichite group minerals), Fe-sulfates (rhomboclase, natrojarosite), and Fe- and Ti-oxides (anatase, nanophase Fe-oxides). Trace amounts of montmorillonite and the general lack of Al-phyllosilicates (except in sample Pfum 3) indicate this area has sustained low fluid:rock ratios through time (Berger and Velde, 1992).

3.4.1.3 Nesjavallavir power plant, Hengill Volcanic Complex

Material collected from the Hengill-1 mudpot contained mostly amorphous SiO2, quartz,

anatase, elemental sulfur, halotrichite group minerals, gypsum, hexahydrite, and montmorillonite,

with moderate amounts of jarosite and alunite, and traces of goethite (Table 13). Jarosite is

consistent with the observed pH range of 2.5 – 3 (Stoffregen, 1993), and Na+/K+ concentrations

favoring Na substitution with [Na+] measured to be an order of magnitude higher than [K+] (Table

11). Gypsum and hexahydrite are present as surficial evaporite crusts, forming from Ca and Mg-

rich fluid as it reaches the surface – Ca originating from pyroxene and feldspar dissolution, and

Mg from olivine and pyroxene dissolution in the parent basalts below. These highly mobile

elements are carried to the surface with the hydrothermal fluid, where they precipitate and form

evaporite crusts. The presence of minor (9 wt %) njarosite and trace (< 1 wt %) nanocrystalline

goethite (nanocrystallinity determined by the presence of broad, diffuse peaks in XRD patterns) in

sample I16 H1 orange, combined with lack of hematite in all Hengill-1 samples indicates at least

intermittent periods of low (~ 2) pH conditions favoring jarosite formation over goethite (Tosca et

al., 2008). Goethite found at Hengill is likely recently formed, and has not yet evolved to the more

stable hematite (Langmuir, 1971) . Observed pH values at the Hengill-1 site range from 2.5 – 3,

indicating that goethite at this site is currently out of equilibrium and may transition to a more

hematite-dominated deposit with time (Langmuir, 1971; Torrent et al., 1982; Tosca et al., 2008).

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The aforementioned gypsum likely formed since this period of low pH, since it would be unstable at pH < 2 and would not form or be preserved in those environments (Klimchouk, 1996).

The Hengill-2 site is a geochemically complex system consisting of three streams of varying pH, temperature, and mineralogical assemblages (Figure 24). The white stream emanates from a cool (22.4C) pH 5 spring without significant contributions from the geothermal field, quickly cooling further to ambient temperature (8.2C) and coating the basaltic talus with a white biofilm rind containing elemental sulfur (H2 white 2) (Figure 25a). Strongly oxidizing conditions

(ORP = 217 mV) at this location suggest that elemental sulfur is forming through H2S oxidation as the fluid meets the surface (Machel et al., 1995; Steudel, 1996). Eh-pH diagrams show a

0 potential pathway for H2S oxidation at this site (Figure 26), with the formation of S as the reduced spring fluid encounters oxidizing conditions at the surface and progresses along the geochemical pathway to sulfate production. The presence of such a large amount of S0 in the white stream indicates this location is out of equilibrium with the geochemical conditions, and there may be kinetic or biologic inhibitors to sulfate production. Determining whether there are microbial influences on sulfur speciation in this stream will have important implications for Martian astrobiology. However, this requires further geochemical investigation of the spring source and the application of gene sequencing techniques to identify the microbial residents and their metabolic pathways at this location.

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Figure 25: Streambed mineralogy at Hengill-2.

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Samples collected from within the gray stream (H2 gray 2) contained coatings of amorphous SiO2, anatase, elemental sulfur, alunite, jarosite, and pyrite (Figure 25a). The pH of the gray stream is not low enough to facilitate jarosite formation (pH < 4). Therefore, the presence of jarosite is likely due to the active downstream transport from the fumaroles located just above the sampling location. The altered material within the gray stream blankets surfaces, appearing to be a precipitated mineral coating rather than a leaching rind. This suggests the Si, Al, Fe, and Ti required were initially mobilized from the underlying parent basalt at pH << 3 (Banin et al., 1997;

Hurowitz et al., 2006; Jercinovic et al., 1990; Paul and Zaman, 1978; Thorseth et al., 1991), and subsequently re-precipitated at the surface, as the pH increased to > 5 (Banin et al., 1997; Crovisier et al., 2003; Drever, 1997; Hurowitz et al., 2006; Paul and Zaman, 1978; Stumm and Morgan,

1996; Thorseth et al., 1991) – forming amorphous SiO2, alunite, anatase, and pyrite. Below the confluence of the white and gray streams (BMH H3), amorphous SiO2 and anatase are no longer precipitating. Pyrite, alunite, and elemental sulfur are still present, with the addition hematite and

Al-phyllosilicates (Figure 25b).

As expected, mineralogy in the red stream is Fe-dominated (Figure 25b), with the bulk of the collected samples (H2 red and BMH H2) consisting of goethite. There is also abundant Al- phyllosilicate (likely kaolinite), and trace amounts amorphous SiO2. A strong oxidation front is evidenced by the presence of pyrite just below a thin crust of kaolinite, anatase, sulfur, and amorphous SiO2 (H2 white) in the hot (88.4C) ground immediately adjacent to the red stream

(Figure 24f).

Hot ground immediately adjacent to the red stream consisted of visually massive multicolored clay-like material (Figure 24f), and is distinctly different from the coated basaltic talus in the gray and white streams. Small (cm-scale) vents dot the surface, precipitating patches

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of elemental sulfur, while the massive material consists primarily of goethite, hematite, pyrite, anatase, alunogen, quartz, amorphous SiO2, and kaolinite, with trace amounts of montmorillonite.

The presence of kaolinite confirms that this region has been extensively altered at low pH and high temperatures (Berger and Velde, 1992; Pirajno, 2010), with patches of montmorillonite indicating regions that are slightly less altered than the kaolinite-rich locations (Berger and Velde, 1992), likely due to localized areas of lower fluid:rock ratios or flow rates. A strong oxidation front is present, with reducing conditions indicated by pyrite located within a centimeter of the surface, and goethite, and hematite above. Present-day ground temperatures in pyrite-rich locations are up to 88.4C, allowing for the possibility that the Fe-sulfides seen at Hengill-2 are due to biological sulfate reduction, which may occur up to ~ 120C (Machel et al., 1995). However, as with the possible geochemical processes in the white stream, this process may also occur through abiotic

H2S oxidation (Rickard, 1997; Rickard and Luther, 1997), and further geochemical and biological investigation is needed to distinguish between these two processes at Hengill.

3.4.2 Implications for Mars hydrothermal deposits

3.4.2.1 Comparison to Mars mineralogy

Alteration products at our Icelandic analog sites were dominated by SiO2, Ca/Al/Fe/Mg- sulfates, Fe/Ti-oxides/hydroxides, and phyllosilicates – similar to assemblages observed on Mars both in situ (Morris et al., 2008; Rice et al., 2010; Ruff et al., 2011; Schmidt et al., 2009) and from orbit (Ehlmann et al., 2016; Skok et al., 2010; Thollot et al., 2012; Weitz et al., 2013, 2011). The lack of abundant amorphous silica (only present in trace amounts) at the high fluid:rock

Þeistareykir mudpots is particularly unusual and suggests areas that do not appear to have abundant silica, such as the western side of Home Plate (Schmidt et al., 2009), may still have formed within a high fluid:rock hot spring system. The other notable lack of mineralogy is the distinct lack of

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alunite group minerals at the Námafjall mudpots. Despite the presence of alunite group minerals at Námafjall fumarole vents, and the presence of other Fe- and Al-sulfates (halotrichite group, rhomboclase, and voltaite) at Námafjall mudpots, no alunite group minerals were identified at the mudpots at this location. Alunite group minerals are common indicators for acidic, often hydrothermal, alteration on Mars (Ehlmann et al., 2016; Farrand et al., 2009; Klingelhöfer et al.,

2004; McCubbin et al., 2009; McHenry et al., 2011; Pirajno, 2010); however, mineralogical analysis of our Námafjall mudpot samples indicates that these diagnostic minerals may not always be present in alteration assemblages, and areas with other Al/Fe-sulfates, Fe-oxides/hydroxides, and phyllosilicates should still be considered as indicators of potential hydrothermal activity.

3.4.2.2 Fe-rich natroalunite detections

The alunite-jarosite solid solution series contains minerals with the formula

+ + 2+ + AB3(SO4)2(OH)6, with Na or K , and occasionally Ca (minamiite), or H3O (hydronium jarosite)

3+ 3+ filling the A site, and Al or Fe filling the B site. KAl3(SO4)2(OH)6 represents the pure alunite endmember and KFe3(SO4)2(OH)6 for pure endmember jarosite. Natroalunite and natrojarosite contain Na+ in the A site. Although minerals spanning the entire alunite-jarosite solid solution series have been made in laboratory settings, it has been rare to find naturally-occurring intermediate members (with a mixture of Fe and Al occupying the B site) of this series (McCollom et al., 2013b). Until McCollom et al. (2013a) identified Fe-rich natroalunite at multiple

Nicaraguan volcanoes, and Black and Hynek (2017) and Beckerman (2016) reported up to ~ 30%

Fe in natroalunite from Costa Rica, the only documented instances of alunite or jarosite with moderate levels of Fe/Al substitution were of a 13% Fe alunite by Brophy et al. (1962) and an 8%

Fe alunite by Alpers et al. (1992). All other documented alunite/jarosite had either no reported Fe levels, or had Fe < 5% or > 80% (Stoffregen and Alpers, 1992; McCollom et al., 2013). Our

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identification of Fe-rich (~60% Fe) natrojarosite at Námafjall in addition to the previously identified Fe-rich natroalunite at other commonly-utilized Mars analog sites, Cerro Negro, Telica, and Masaya volcanoes in Nicaragua (Hynek et al., 2013; Marcucci et al., 2013; McCollom et al.,

2013a) and Poás and Turrialba in Costa Rica (Beckerman, 2016; Black and Hynek, 2017) suggest that Fe/Al substitution in the alunite-jarosite subgroup may be a common process in hydrothermal systems, resulting in intermediate members (15 – 80% Fe in the B site) of this solid solution series.

As demonstrated by McCollom et al. (2014), Mössbauer spectra of jarosite and Fe-rich natroalunite are indistinguishable, and other methods such as VNIR, XRD, Raman, or EDS must be used to characterize the Fe amounts that are present in minerals of this subgroup. McCollom et al. (2013b) suggest that Fe-rich natroalunite may be an initial and thermodynamically unstable phase which forms in the early stages of alteration and later alters to endmember alunite/jarosite and associated

Fe-oxide phases such as hematite. Our results from Námafjall are consistent with this hypothesis, as the Fe-rich natroalunite (sample Nama U4) is found with amorphous SiO2, natroalunite, and gypsum (Table 10, Figure 18) – the same initial alteration assemblage observed by McCollom et al. (2013b). All instances of (natro)alunite in our samples are found in association with Fe- oxides/hydroxides and/or pyrite (Table 10Table 12Table 13) except for one (Nama U8 Grad 1), as expected from McCollom et al.’s (2013b) hypothesis. Nama U8 Grad 1 contains only natroalunite

(Figure 18). The lack of Fe-oxides/hydroxides/sulfides suggests this sample did not evolve from a more Fe-rich alunite-jarosite phase. Additionally, all instances of (natro)jarosite are found with

Fe-oxides/hydroxides and/or pyrite as well as Al-sulfates or phyllosilicates (Table 10Table

12Table 13), indicating a possible source for additional Fe to transform from an intermediate phase alunite-jarosite to endmember jarosite, and a secondary phase for Al that is replaced and removed from the crystal structure of the intermediate composition (Al-rich) jarosite. Additional high

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resolution investigation would be needed for further insight into the mineral associations in these samples, but is beyond the scope of this study. However, the identification of another hydrothermal system with naturally-occurring Fe-rich natroalunite indicates these intermediate alunite-jarosite products may be more common than previously thought, and must be considered as possible secondary minerals on Mars (McCollom et al., 2013; 2014). The geochemical conditions necessary for formation, and even the detailed crystal structure of intermediate alunite- jarosite products are not yet understood (Stoffregen and Alpers, 1992; Desborough et al., 2010;

McCollom et al., 2013) and would benefit from detailed laboratory investigations to determine the thermodynamic and physical properties of these minerals. The combination of distinct alunite and jarosite peaks in our Raman spectra from sample Nama U4 (a mixture of natroalunite and 61% Fe natroalunite per bulk XRD) (Figure 27), rather than the gradual shift that is observed by McCollom et al. (2014) in their synthetic intermediate alunite-jarosite materials, is suggestive of varied composition alunite and jarosite that is beyond the resolution (< 2 m for Raman) of our instrumentation. Indeed, investigation of this sample with SEM-EDS suggests this intermediate chemistry is due to a chemically-variable mixture of natroalunite-natrojarosite with crystals ranging from < 2 m up to 10 – 15 m in diameter (Figure 27). A better understanding of the geochemical processes and conditions which result in intermediate members of the alunite-jarosite series may help to further constrain the geochemical history of both terrestrial and martian hydrothermal sites in the future.

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3.5 Conclusions

Our ability to identify and understand the geochemical history of hydrothermal deposits on

Mars is dependent on our knowledge of their terrestrial counterparts. For this study, we characterized the bulk surface mineralogy of three Icelandic hydrothermal sites using VNIR reflectance spectroscopy and bulk powder XRD, including three previously uncharacterized mudpots and a fumarole at the Þeistareykir high temperature geothermal field in Iceland’s

Northern Volcanic Zone. In addition to Þeistareykir, two locations at the Nesjavallavir Geothermal

Plant on Hengill’s northern slopes, and several mudpots and fumaroles at the Námafjall geothermal

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field were also characterized. We focused on surface expressions of hydrothermal systems rather than deep alteration mineralogy as surface deposits are more likely to be exposed and accessible on the surface of Mars. The thin present-day Martian atmosphere and absence of persistent surface water should result in minimal erosion rates since the end of the Hesperian (Ehlmann et al., 2011b;

Fassett and Head, 2008; Hoke et al., 2011; Hynek et al., 2010) and the majority of Mars’ volcanic activity (Robbins et al., 2011), thus preserving surficial hydrothermal deposits. Alteration assemblages at our sampling locations included a variety of Al/Ca/K/Fe/Mg-sulfates, amorphous and crystalline SiO2, Fe/Ti-oxides and hydroxides, and phyllosilicates.

One sample contained > 20 wt % of an intermediate member of the alunite-jarosite solid solution series (~60% Fe in the B site as identified by XRD). Naturally-occurring intermediate members (15 – 80 % Fe) of this solid solution series have only been reported in Cerro Negro,

Telica, and Masaya Volcanoes in Nicaragua (Hynek et al., 2013; Marcucci et al., 2013; McCollom et al., 2013a; McCollom et al., 2014), Poás and Turrialba Volcanoes in Costa Rica (Beckerman,

2016; Black and Hynek, 2017). Our identification of ~60% Fe natroalunite at Námafjall suggests intermediate members of this solid solution series may be common hydrothermal alteration products. Investigation of this sample with Raman laser spectroscopy yielded spectra with both distinct alunite and jarosite peaks, suggesting that these intermediate products may actually be finely layered or microcrystalline alunite-jarosite that are only discernible with SEM (Figure 27).

However, additional investigation is needed to better understand the physical form and geochemical conditions formation for these intermediate products.

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4. Hydrothermal acid-sulfate alteration of a synthetic Mars composition basalt: Implications for hydrothermal deposits on Mars

Note: This chapter has been prepared to submit to the Journal of Geophysical Research: Planets as: Black, S. R., Hynek, B. M., Larsen, J. F., and McCollom, T. M. “Hydrothermal acid-sulfate alteration of a synthetic Mars composition basalt: Implications for hydrothermal deposits on Mars” Journal of Geophysical Research: Planets. All work was completed by SRB. JFL directed and assisted with Mars basalt creation at the University of Alaska Fairbanks.

Abstract: Experimental acid-sulfate alteration of both glassy and crystalline synthetic Mars basalts (17.0 and 19.8 wt % FeOT, respectively) resulted in closed-system alteration assemblages dominated by Fe-sulfates (Fe-bearing alunite group, rhomboclase, and voltaite), and Fe- oxides/hydroxides. Closed-system alteration of terrestrial parent materials spanning the felsic to mafic range also produced intermediate alunite group sulfates, but lacked additional Fe-bearing secondary mineralogy, despite the supersaturation of hematite, goethite, and nontronite in several instances, suggesting the presence of kinetic barriers to their formation in our experimental setups. Open-system alteration of our synthetic Mars basalts produced an initial alteration assemblage dominated by Fe-sulfates (alunite group, rhomboclase, and voltaite), but quickly transitioned to SiO2, Al-sulfate, and Ti-oxide dominated as more soluble cations were leached from the system. Fe-sulfates continued to form in the dried gel phase, suggesting their incorporation into the rapidly precipitating amorphous SiO2 coating as Fe was leached from the substrate. Combined projections from our idealized closed system results and previous terrestrial analog investigations, we expect ~15 – 35 wt % total Fe-bearing secondary mineralogy to result from acid-sulfate alteration of Martian basalts – similar to abundances we see in high-Fe terrestrial analog systems.

4.1. Introduction

4.1.1. Hydrothermal systems on Mars

A long history of volcanic activity on Mars has been supported by both remote and in situ observations of volcano structures and deposits (Crumpler et al., 1996; Robbins et al., 2011;

Squyres et al., 2007; e.g. Wilson and Head, 1994). In addition to volcanism, surface and subsurface water appears to have played a large role in Mars’ geologic history (Ehlmann and Edwards, 2014;

Hynek et al., 2010; Stuurman et al., 2016; Wray et al., 2009). The combination of volcanic activity and abundant crustal water suggests hydrothermal systems were common on early Mars, and may have persisted through much of Mars’ geologic history (Hynek et al., 2013). Enriched sulfur levels

(average ~6 wt %) in the Martian basaltic crust indicate high volumes of SO2 degassing from

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Martian volcanic systems (King and McLennan, 2010; Peretyazhko et al., 2018; Righter et al.,

2009), making it likely that Mars hosted abundant acid-sulfate hydrothermal systems.

The mineralogical signatures of acid-sulfate hydrothermal systems have been identified across the surface of Mars. The Home Plate deposit within Gusev crater has been interpreted as an altered pyroclastic material, and contains large amounts of SiO2, hematite, and Fe/Mg/Ca-sulfates

(Morris et al., 2008; Ruff, 2015; Ruff et al., 2011; Squyres et al., 2008, 2007; Yen et al., 2008).

Orbital observations have also identified SiO2 deposits within Nili Patera (Skok et al., 2010), a

Yellowstone-scale area of the common hydrothermal mineral alunite within Cross crater (Ehlmann et al., 2016), and mixed sulfates and phyllosilicates in parts of Noctis Labyrinthus (Thollot et al.,

2012; Weitz et al., 2011).

Our understanding of the environmental conditions which formed these deposits, and in turn their potential for habitability, is dependent on our knowledge of the geochemical processes that occur within a hydrothermal system. Geochemical conditions (e.g. pH, temperature, fluid:rock) are interpreted through the lens of secondary mineralogy. To develop a thorough picture of the geochemical history of any Martian hydrothermal system, we must better understand the influence which primary basalt composition has on secondary alteration mineralogy.

4.1.2. Martian basalt compositions

Basalt compositions on Mars have been characterized by a number of methods. Orbital instruments such as the Compact Reconnaissance Imaging Spectrometer for Mars (CRISM)

(Murchie et al., 2009b), Thermal Emission Spectrometer (TES) (Christensen et al., 2001), and the

Gamma Ray Spectrometer (GRS) (Boynton et al., 2004) allow for general estimates of basalt composition across large regions of Mars, while direct measurements of SNC meteorites (Lodders,

1998) and in situ rover observations (McSween et al., 2006; Ming et al., 2008) result in more

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detailed compositional data in select locations. Due to the limited number of landed missions and

SNC meteorites, the majority of our understanding of Martian basalt composition relies on orbital data such as Thermal Infrared (TIR) (Christensen et al., 2005; 2001) emission and Visible Near-

Infrared (VNIR) reflectance spectroscopy (Murchie et al., 2009b; Rogers and Hamilton, 2015).

Orbital TIR and VNIR spectra show four general basalt types, characteristic regions of which are the southern highlands consisting of primarily plagioclase and clinopyroxenes, Noachian regions enriched in olivine and low-calcium pyroxene (LCP), Hesperian lava flows with lower abundances of olivine and LCP, and the northern plains, which have the lowest pyroxene abundances and increased abundance of silica phases (Mustard et al., 2005).

In situ measurements by the Mars Exploration Rovers (MER), Spirit and Opportunity, as well as those by the Mars Science Laboratory (MSL) rover, Curiosity, also provide elemental abundances through the use of the Alpha Particle X-Ray Spectrometer (APXS) instruments that are included in the science payloads (e.g. Rieder et al., 2003). The Pathfinder lander also included an APXS instrument (Rieder et al., 1997) but no drill or brush was included on the lander, so all

APXS measurements were of surficial weathering rinds and not the primary basalts (McSween et al., 2009; Michalski et al., 2005; Wyatt et al., 2004). Pathfinder APXS measurements are therefore not included in this study. With the inclusion of the Rock Abrasion Tool (RAT) or drill on both the MER and MSL rovers (Gorevan et al., 2003; Okon, 2010), the APXS instruments are able to sample below dust and weathering rinds and provide our most detailed look at Martian basalt compositions. In situ and laboratory analyses show significant Fe enrichment in Martian basalts

(up to ~21 wt % FeOT in Gusev crater and Shergottite LA 1) relative to terrestrial samples (Table

14).

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Table 14: Elemental compositions of unweathered Martian basalts and commonly-utilized terrestrial analogs

Sample SiO2 FeOT Al2O3 MgO CaO Na2O K2O MnO TiO2 Total Martian basalts Adirondack, Gusev 1 45.3 21.09 10.43 11.9 7.76 2.09 0.03 0.42 0.49 99.51 Backstay, Gusev 2 49.4 13.1 13.1 8.3 6.0 4.0 1.02 0.25 0.93 96.10 Irvine, Gusev 2 47.5 19.7 8.3 9.5 5.8 3.0 0.60 0.37 1.06 95.83 Algonquin, Gusev 2 41.9 20.9 6.4 16.0 4.0 2.3 0.40 0.39 0.58 92.87 Bounce Rock, Meridiani 3 50.8 15.6 10.1 6.4 12.5 1.3 0.10 0.43 0.78 98.01 Shergottite LA 1 4 49.1 21.2 11.2 3.53 10.0 2.22 0.24 0.45 1.30 99.24 Shergottite QUE 94021 5 47.9 18.5 11.0 6.3 11.4 1.58 0.05 0.45 1.84 99.02 Bost (2012) Representative Mars Basalt 47.58 18.00 10.87 11.90 6.12 2.94 0.19 0.29 0.76 98.65 Terrestrial basalts Hengill, Iceland 6 50.10 13.90 13.40 4.86 9.61 2.84 0.68 0.27 2.89 98.55 Krafla, Iceland 7 48.54 11.58 15.39 8.80 12.71 1.83 0.11 0.18 1.11 100.25 Kilauea, Hawaii – Pu’u O’o 8 52.32 10.77 13.48 6.90 10.99 2.05 0.41 0.17 2.44 99.53 Kilauea, Hawaii – 1974 flow 8 51.87 10.59 13.30 7.44 11.02 2.71 0.47 0.17 2.48 100.05 Cerro Negro, Nicaragua 9 49.73 9.70 19.50 4.73 11.50 2.18 0.43 n.r. n.r. 97.77 Telica, Nicaragua 9 51.72 9.60 19.10 4.13 10.10 2.84 0.96 n.r. n.r. 98.45 Momotombo, Nicaragua 9 54.40 9.10 16.90 4.47 9.20 2.89 0.91 n.r. n.r. 97.87 Poás, Costa Rica 10 53.30 8.12 17.46 4.36 8.49 2.95 1.08 n.d. 0.78 96.54 Turrialba, Costa Rica 10 54.14 7.59 16.30 6.02 9.22 3.12 1.43 0.01 1.05 98.88 Mauna Kea, Hawaii 11 49.74 6.02 17.37 3.93 6.60 4.33 1.90 0.21 2.77 92.87 n.d. = not detected n.r. = not reported 1 back-calculated end member composition from McSween et al. (2006) 2 Ming et al. (2008) 3 Rieder et al. (2004) 4 Rubin et al. (2000); Hurowitz et al. (2005) 5 Lodders (1998) 6 Trønnes (1990) 7 Nicholson and Latin (1992) 8 Chemtob and Rossman (2014) 9 Hynek et al. (2013) 10 Black and Hynek (2017) 11 Morris et al. (2000)

4.1.3. Previous work

Experimental studies of acid-sulfate alteration have been completed a number of times, with a wide variety of conditions and parent materials (Table 15). While many experimental setups have made use of terrestrial parent materials with lower parent FeOT (Banin et al., 1997; Golden et al., 2005; Hausrath and Tschauner, 2013; Marcucci and Hynek, 2014; McCollom et al., 2013b), acid-sulfate alteration of high-Fe terrestrial basalt or synthetic “Mars” basalt has also been conducted at both high and low temperatures (Baker et al., 2000; Horgan et al., 2017; Hurowitz et al., 2005; Peretyazhko et al., 2018; Tosca et al., 2004; Yant et al., 2016). Baker et al.’s (2000) alteration of a 17.65 wt % FeOT Columbia River Basalt focused on carbonate formation and 128

secondary mineralogy in Martian meteorites using a CO2-saturated hydrothermal fluid/vapor at temperatures up to 400ºC. Secondary mineral assemblages contained various SiO2 phases, carbonates, oxides, and occasionally zeolites and secondary silicates. Tosca et al. (2004) revisited

Banin et al.’s (1997) acid fog alteration experiments, which were originally run with Hawaiian tephra (11.79 wt % FeOT), and re-ran the alteration experiments using synthetic Mars Pathfinder rock and soil (13.73 and 19.23 wt % FeOT, respectively). Alteration of these synthetic Pathfinder rocks and soils using an H2SO4 + HCl fluid mixture produced abundant Mg/Fe/Ca/Al/Na-sulfates, halite, Fe-oxides and phosphates, and amorphous SiO2, and Yant et al.’s (2016) return to this experiment with a focus on spectral characteristics of the resulting alteration products yielded similar results. Banin et al.’s (1997) original acid fog alteration experiments with lower FeOT

Hawaiian tephra resulted in only Ca/Al-sulfates, with the lack of Fe/Mg-bearing phases interpreted as formation of poorly or nanocrystalline phases that could not be detected by XRD. Tosca et al.’s

(2004) results indicated that bulk chemical composition (such as wt % FeOT) and mineralogy of the parent material (i.e. what phases the various elements are partitioned into and their thermodynamic properties) are primary controls on secondary alteration mineralogy. These factors have a significant influence on hydrothermal fluid composition, and therefore, secondary mineral formation. Subsequent alteration experiments using a synthetic LA Shergottite by Hurowitz et al.

(2005) found the dissolution of plagioclase to be a dominant processes, although their primary basalt did not contain any detectable olivine, which should also influence fluid chemistry with a rapid release of Fe and Mg to the alteration fluid (Jonckbloedt, 1998). Most recently, Peretyazhko et al. (2018) produced secondary smectite through high-temperature (200ºC) alteration of a synthetic Adirondack-type glassy Mars basalt, and concluded that smectite type was dependent primarily on Mg release from the parent material.

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Here, we conducted the first long-duration (70 – 140 days) laboratory-based acid-sulfate alteration experiments using both glassy and crystalline olivine-bearing synthetic Mars basalts

(17.0 and 19.8 wt % FeOT, respectively). We also altered terrestrial parent materials from basalts to rhyolite/obsidian, to systematically investigate the effect of parent FeOT on secondary mineralogy. Experiments were conducted as both open and closed systems, with both crystalline and glassy synthetic Mars basalts, to further investigate the influence of parent mineralogy and open vs. closed systems.

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Table 15: Previous laboratory-based acidic alteration studies Study Parent material Conditions Secondary mineralogy

Banin et al. (1997) Partly palagonitized Hawaiian tephra H2SO4, HCl, HNO3 Alunogen, gypsum, kieserite, pH 2.59 – 7.74, epsomite, nanophase Fe/Mg-bearing 25C, 7 days minerals possible

Baker et al. (2000) Columbia River 17 wt % Fe basalt CO2 + H2O Opal-CT, hydrated SiO2, quartz, 23, 75, 200, and 400C calcite, siderite, ankerite, magnesite, 200:1, 1:1, and 1:10 fluid:rock hematite, maghemite, sacrofanite, 4 – 7 days vesuvianite, sepiolite, richterite, biotite

Tosca et al. (2004) Simulated Pathfinder basalt and soil H2SO4 + HCl Hexahydrite, rhomboclase, 25C, 14 days melanterite, general Fe-sulfates, phosphates, and oxides, gypsum/anhydrite, alunogen, tamarugite, halite, amorphous SiO2

Golden et al. (2005) Hawaii sand and tephra Closed system: Amorphous SiO2, gypsum/anhydrite, H2SO4 (vapor and immersed) hexahydrite, alunogen, voltaite, 145C, 6 days jarosite, halite Flow through: H2SO4 + H2O2 145C, three steps of 48 hours

Hurowitz et al. (2005) Simulated LA Shergottite H2SO4 + HCl (pH 1.1 and 3.6) Amorphous SiO2, lepidocrocite, 76C, 14 – 17 days ferrihydrite, anhydrite, alunogen, 100:1 and 1000:1 fluid:rock ferrinatrite, rhomboclase, Ca- and Na- flow through and immersed sulfates

Hausrath et al. (2013) Olivine, fluorapatite, basaltic glass H2SO4 + H2O vapor Amorphous SiO2, anhydrite, kieserite, 155C + ambient hexahydrite, Ca-, Al-, and Mg-sulfates, 6, 24, and 72 hours Ca-phosphates

McCollom et al. (2013b) Cerro Negro basalt cinders 1M H2SO4, 145C Amorphous SiO2, anhydrite, Fe-oxide, ~ 1:1 to 10:1 fluid:rock natroalunite, alunogen, hexahydrite 7 – 137 days tamarugite, kieserite, unidentified Fe- silicate and Mg-, Al-, Fe-, and Ca- sulfates

Marcucci and Hynek (2014) Cerro Negro basalt cinders, 1M H2SO4 Amorphous SiO2, gypsum/anhydrite, plagioclase, pyroxene, olivine 65, 15, and 200C ferrohexahydrite, hexahydrite, 1:1, 4:1, and 10:1 fluid:rock pentahydrate, starkeyite, natroalunite, 3 days – 7.5 months alunogen, voltaite, rhomboclase, kieserite, Fe-oxide, Ca-, Al-, Mg-, and Fe-sulfates

Yant et al. (2016) Simulated Pathfinder basalt H2SO4 + HCl (pH 0 – 4) Amorphous SiO2, gypsum/anhydrite, (from Tosca et al., 2004) Ambient temperature alunogen, pickeringite, halotrichite, 1:1 fluid: rock, 14 days hexahydrite, voltaite, melanterite, rozenite, Na-Alum, tamarugite, ilesite, blödite, szomolnokite

Horgan et al. (2017) Columbia River, Icelandic, & H2SO4 (pH 1 – 3) Amorphous SiO2, gypsum, Al/Mg- Hawaiian basalts 25C; 5:1 fluid:rock sulfates (alunite, kieserite), oxide 213 days; flow through coating, unidentified high-SiO2 phase

Peretyazhko et al. (2018) Glass-rich synthetic Mars 0.011 – 0.0425 M H2SO4 Hematite, anhydrite, natroalunite, (Adirondack) basalt pH ≤ 2 dioctahedral smectite, saponite 200C; 60:1 fluid:rock 14 days

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4.2. Methods

4.2.1. Samples

4.2.1.1. Synthetic Mars basalt (SMB)

Synthetic Mars basalts (herein referred to as SMB) were created from pure mixed oxides

(SiO2, TiO2, Al2O3, Fe2O3, MgO, and MnO) and carbonates (CaCO3, Na2CO3, and K2CO3) at the

University of Alaska Fairbanks experimental petrology lab using a 1 atmosphere DelTech gas-

mixing furnace. The Bost (2012) average Mars basalt composition (Table 14) was used as a

starting point when mixing pure oxides and carbonates. Additional Fe absorption by the Pt crucible

was expected during the creation of the crystalline sample due to the extended time spent at high

temperatures (Grove, 1981), therefore, additional Fe2O3 was added to the powder mixture for the

crystalline basalt. Due to the completion of several crucible-conditioning runs before sample

creation the observed Fe absorption was minimal, yielding FeOT abundances up to an Irvine-type

19.8 wt % (Table 16). Oxygen fugacity (fO2) calculations of spinel in several SNC meteorites

yields fO2 values between the Iron-Wustite (IW) and Fayalite-Magnetite-Quartz (QFM) mineral

buffers (Herd et al., 2001), therefore, our fO2 for reduced basalt formation was set at IW + 1.73

(+/- 0.12) log units, with a gas flow rate of ~130 mL/min CO2 + H2.

Table 16: Elemental compositions of parent materials measured by XRF analysis

Sample SiO2 FeOT Al2O3 MgO CaO Na2O K2O MnO TiO2 Total Synthetic Mars Basalt: Crystalline SMBxl 46.5 19.8 10.9 11.3 6.1 2.4 0.2 0.3 0.8 98.3 Synthetic Mars Basalt: Glassy SMBgl 48.2 17.0 11.4 11.5 6.3 2.6 0.2 0.3 0.9 98.4 Hengill basalt HN-BAS 46.5 11.8 14.4 6.6 12.2 1.9 0.2 0.2 2.2 96.0 Cerro Negro basalt CN-BAS 49.8 10.1 18.0 6.1 11.3 2.1 0.5 0.2 0.8 98.9 Turrialba basaltic andesite B-AND 54.1 7.6 16.3 6.0 9.2 3.1 1.4 < 0.1 1.1 98.8 Flagstaff andesite AND 69.3 2.5 16.5 1.0 0.4 0.4 5.5 < 0.1 0.3 95.9 Ruby Mt. rhyolite RHY 75.8 0.5 13.0 0.4 0.3 3.7 4.6 0.1 0.1 98.5 Millard County obsidian OBS 76.4 0.9 13.1 0.1 0.8 4.0 4.8 0.1 0.1 100.3

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Glass-dominated samples (SMBgl) were run in an Fe-conditioned Pt crucible at 1330 (+/-

4)C for 4-6 hours and drop quenched into water. Experiments were initially examined using backscattered electron imaging on a JEOL JXA-8530 Superprobe electron probe microanalyzer

(EPMA) with field emission column. Mineral phase identification employed use of the attached

Thermo Noran System 7 energy-dispersive spectrometer (EDS). Glass composition and sample homogeneity were initially confirmed by EPMA-EDS analysis, with additional composition quantification using Wavelength-Dispersive Spectrometry (WDS). Glass-dominated samples are

75 or more volume % glass, with up to 25 volume % Mg/Fe spinel and euhedral Fo79 olivine – indicative of melt conditions at or near equilibrium (Table 17, Figure 28).

Crystalline basalts (SMBxl) were created using the same initial oxide/carbonate powder mixture. The crystallization run had a total duration of 76.5 hours, with the experiment held at the initial 1330 (+/- 4)C temperature for 4 hours and then ramped down by 6C per minute to 959C.

Oxygen fugacity was held at IW +1.75 (+/-0.39) log units. XRD, EPMA-WDS, and petrographic analysis of products found 21 wt % Fo61 olivine, 39 wt % An52 plagioclase feldspar (labradorite),

23 wt % pyroxene (augite), and 17 wt % glass (Table 17, Figure 28).

Table 17: Mineral components of parent rocks as determined by XRD, EPMA, and/or petrographic thin section

Source Wt % FeOT Analysis Major components Minor components XRD, EPMA, An labradorite (39 wt %), augite (23 wt %), Basaltic glass, SMBxl Synthetic 19.8 52 thin section Fo61 olivine (21 wt %) titanomagnetite Basaltic glass (~75 vol %), SMBgl Synthetic 17.0 XRD, EPMA Mg/Fe spinel (~1 vol %) Fo79 olivine (~24 vol %), XRD, EPMA, An bytownite (40 wt %), augite (30 wt %), Secondary chlorite HN-BAS Hengill, Iceland 11.8 80 thin section Fo79 olivine (8 wt %), basaltic glass (18 wt %) (8 wt %) XRD, EPMA, An bytownite (61 wt %), diopside (18 wt %), CN-BAS Cerro Negro, Nicaragua 10.1 83 thin section Fo70 olivine (8 wt %), basaltic glass (13 wt %) XRD, Labradorite (57 wt %), diopside (23 wt %), B-AND Turriabla, Costa Rica 7.6 Olivine (6 wt %) thin section glass (15 wt %) XRD, Quartz (63% wt %), orthoclase (26%), AND Flagstaff sill, Boulder, CO 2.5 Biotite (5 wt %) thin section glass (7 wt %) XRD, Quartz (66 wt %), sanidine (22 wt %), Biotite (< 1 wt %), RHY Ruby Mountain, CO 0.5 thin section glass (8 wt %) Secondary chlorite (4 wt %)

OBS Millard County, UT 0.9 XRD SiO2 glass (~100 wt %)

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4.2.1.2. Terrestrial parent materials

High Fe basalts from Hengill volcano (HN-BAS) were collected from the Nesjavallavir

Power Plant in the northern portion of the Hengill region in August 2016. Basalts collected from

this location contain phenocrysts of plagioclase feldspar (bytownite, An80), clinopyroxene

(diopside), and olivine (Fo79). The fine matrix is composed of glass, plagioclase feldspar, and

clinopyroxene microlites (Table 17, Figure 28). In addition to high Fe basalts from Hengill, we

chose to include a second basalt with lower FeOT (Table 16) from the commonly-utilized Mars

analog, Cerro Negro (CN-BAS). The CN-BAS parent rock was collected from the 1992 crater in

2013 and contains large vesicles and abundant phenocrysts of olivine (Fo70), pyroxene (diopside),

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and plagioclase (An83), with a groundmass of plagioclase, pyroxene, and glass (Table 17, Figure

28).

Fresh basaltic andesites (B-AND) were collected from the summit crater of Turrialba

Volcano in Costa Rica during a November 2013 field campaign. Secondary mineralogy from in situ acid-sulfate alteration of these basaltic andesites are reported in Black and Hynek (2017).

Petrographic thin sections show large plagioclase phenocrysts, large and relatively unaltered olivines with some slight Fe exsolution rims, and a groundmass composed primarily of pyroxene microlites with a small amount of glass (Table 17, Figure 28). A biotite-rich andesite (AND) from the Flagstaff Sill in Boulder, Colorado (Hoblitt and Larson, 1975) was also included as an intermediate parent rock. Large biotite crystals are visible both in hand sample and thin section

(Table 17, Figure 28) in addition to quartz and alkali feldspar. Despite the mineralogical similarity to our rhyolite parent sample, the overall abundance of biotite within the AND parent rock sets the bulk rock FeOT and MgO abundances apart from our rhyolite parent (Table 16).

The identification of felsic trachyte clasts by MSL Curiosity at Gale crater (Sautter et al.,

2016) and orbital CRISM identifications of quartz and feldspar within Syrtis Major (Smith and

Bandfield, 2012) suggest there are more evolved igneous materials present on Mars in addition to the abundant mafic materials. Therefore, our experimental setup also includes felsic parent materials rhyolite and obsidian. The Ruby Mountain rhyolite (RHY) (Christiansen et al., 1983) consists of large quartz and alkali feldspar phenocrysts, with small amounts of biotite and a groundmass of feldspar, quartz, glass, and minor secondary chlorite (Table 17, Figure 28).

Obsidian (OBS) collected from Millard County, Utah (Crecraft et al., 1981) was also used as an analog for potential extrusive felsic volcanics on Mars. The obsidian is similar in composition to the RHY parent material, with slightly increased alkalis and FeOT (Table 16), and no apparent

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crystalline material observed in XRD (Table 17). Inclusion of both crystalline and glassy felsic endmembers will allow for comparison of alteration products between these two forms due to their differing thermodynamic properties.

4.2.2. Experimental setup

Closed-system alteration experiments (Table 18) for each parent lithology (SMB, high and low Fe terrestrial basalts, basaltic andesite, andesite, rhyolite, and obsidian) were done at 95C 

5C in a 1 atm oven. Approximately 3.0 g of each parent material was crushed to a mixture of sand to fine gravel-size particles, added to a Savillex teflon reaction vessel, and immersed in 1M

H2SO4 (4:1 fluid:rock by weight) for 70 days to ensure alteration had occurred in even the most felsic parent materials. Upon removal from the oven, each vessel was weighed to confirm no loss of fluids or volatiles occurred during the reaction. The fluid was removed, filtered, and tested for pH (using Micro Essential Labs Hydrion narrow range pH test strips) and major element fluid chemistry. Altered samples were rinsed with ethanol, and the SiO2-rich gel coating separated from the solids to isolate materials formed during the primary reaction phase (solids) from those that formed during subsequent evaporation of the hydrothermal fluids at room temperature (dried gel)

(Figure 29).

Additional long-term open-system alteration experiments were conducted using both the crystalline and glassy SMB (Table 18). For each material (SMBxl and SMBgl), approximately 4 g of parent rock were ground to sand-size particles. As with the closed system experiments, reaction vessels were kept at 95C  5C in a 1M H2SO4 solution (5:1 fluid:rock by weight). To simulate a flow through system, fluid was removed and replaced with new 1M H2SO4, filtered, and tested for pH and major element fluid chemistry every two weeks for 140 days. A small amount (~0.1 g) of altered material was also removed and dried every two weeks for mineralogical

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analysis via bulk powder XRD and SEM-EDS. This material was separated by the previously

described methods. Upon completion of the experiment, fluid was again removed and prepared

for analysis.

Table 18: Experimental conditions for each laboratory set up Parent Mass Temperature Starting fluid Duration Fluid:rock System SMBgl SMBxl HN-BAS CN-BAS 4:1 3 g 95C  5C 1M H SO 70 days Closed B-AND 2 4 by weight AND RHY OBS SMBxl 5:1 4 g 95C  5C 1M H SO 140 days Open SMBgl 2 4 by weight

Figure 29: Parent materials, post-reaction vessels, and dried products from closed system reactions

4.2.3. Sample analysis

4.2.3.1. Fluid chemistry

Fluids were removed, filtered, and saved for analysis at the conclusion of the 70-day

closed-system experiments. For open-system experiments, fluids were removed, filtered, and

saved for analysis every 14 days as well as at the conclusion of the 140-day experiment. Fluid

chemistry was analyzed for major cation and anion concentrations in the Laboratory for

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Environmental and Geological Sciences at CU Boulder. Major cations were measured using an

ARL 3410+ inductively coupled optical emission spectrometer (ICP-OES) with a blank and three

standards for ICP-OES calibration. Anions were measured using a Dionex 4500i ion

chromatograph using five standards for calibration and an AG14 guard column and AS14 column

for separation of anions.

4.2.3.2. X-Ray Diffraction (XRD)

Parent rock, altered solids, and dried gel were analyzed with bulk powder XRD using an

Olympus Terra portable XRD (Olympus Scientific Solutions, 2014; Sarrazin et al., 2005) with a

Cu-Kα radiation source, 5 – 55º 2θ detection window, ~0.05 º 2θ resolution, and ~1 wt % detection

limit (Treiman et al., 2010). All samples were dried at room temperature, then powdered to ~10

µm, and a small amount (<< 1 g) was inserted into a mylar sample cell. All analyses were run to

250 exposures to maximize signal:noise. Diffraction patterns were processed using the Xpowder

software, and compared to the American Mineralogist Crystal Structure Database “difdata” library

(Lafuente et al., 2015) and McCollom et al.’s (2014) alunite-jarosite series diffraction patterns.

4.2.3.3. Visible Near-Infrared (VNIR) Spectroscopy

Final alteration products for both closed and flow through systems were analyzed with a

TerraSpec 4 High Resolution Visible Near-Infrared (VNIR) spectrometer from Malvern

Panalytical (ASD Inc., 2017). The TerraSpec is a contact probe with an internal halogen light

source, which measures reflected light from 350 to 2500 nm, with a 6 nm/band spectral resolution.

Resultant reflectance spectra were processed in the Environment for Visualizing Images (ENVI)

software, and absorptions matched to library spectra from the USGS (Clark et al., 2007), CRISM

tools (PDS Geosciences Node and CRISM Spectral Library Working Group, 2014), and

McCollom et al.’s (2014) alunite-jarosite series.

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4.2.3.4. Scanning Electron Microscope Energy-Dispersive Spectroscopy (SEM-EDS)

Both solid and dried gel alteration products were analyzed with SEM-EDS at the USGS

Spectroscopy Laboratory in Denver, Colorado. Dry materials were imaged and analyzed using the

JEOL JSM-5800LV SEM with a Thermo Noran SDD-EDS and 15 keV accelerating voltage.

Several of the dried gel products were strongly hygroscopic and required the use of the

Environmental (low vacuum at 20 Pa) FEI Quanta 450 FEG-SEM for imaging at 15 keV and EDS

analysis with the Oxford Instruments SDD-EDS using the Aztec software package. EDS does not

allow for quantitative analysis of compositions, however, mineral phases may be identified (Reed,

2005).

4.3. Results

4.3.1. Closed system

4.3.1.1. Crystalline synthetic Mars basalt (SMBxl)

Closed-system alteration of SMBxl parent rock resulted in a mixture of residual primary

pyroxene and feldspar, with an abundance of secondary Mg-sulfates (hexahydrite/epsomite), linear

crystals of anhydrite, large crystalline Fe-oxides/hydroxides (hematite, maghemite, goethite),

amorphous SiO2, and small areas of Fe/Al/K/Na-sulfate that appear to be alunite group minerals

(Figure 30a-c, Table 19). Fluids collected from the reaction vessel contained 5.08 x 105 ppb Fe –

five times higher than Fe concentrations in alteration fluids from HN-BAS (1.05 x 105 ppb), CN-

BAS (1.97 x 105 ppb), and B-AND (1.01 x 105 ppb) (Figure 31), indicating the potential for further

Fe-bearing mineral formation upon complete evaporation of these hydrothermal fluids. Minerals

with a calculated saturation index (SI) > -0.4 (SI = log Q/K, where Q = the reaction quotient and

K = the equilibrium constant), indicating equilibrium (SI = 0) or supersaturated (SI > 0) conditions,

are shown in Table 20. We consider minerals with -0.4 < SI < 0.4 to be at equilibrium to account

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for any errors within the thermodynamic database. The Fe-smectite nontronite is supersaturated with respect to the SMBxl fluid (SI = 1.2916), but has not been identified in the secondary mineral assemblage (Table 19), suggesting kinetic inhibitors to its formation such as an abbreviated maturation time.

Figure 30: A-C) Solid alteration products from the closed-system SMBxl run; D) Solid alteration products from the closed-system SMBgl run; E and F) Dried gel products from the SMBgl separated gel material; G) Solid alteration products from HN-BAS; H and I) Solid alteration products from CN-BAS; J) Solid alteration products from B-AND; K) Solid alteration products from AND; L) Dried gel products from AND; M) Solid products from RHY; N) Solid products from OBS; O) Dried gel products from OBS. Alg = alunogen; Hx hexahydrite; Pk = pickeringite; Fe-ox = Fe-oxide; An = anhydrite; Si = amorphous SiO2; Nat-j = natrojarosite; Fe-alun = Fe- bearing alunite; Nat-a = natroalunite; Fe/Al/Mg/mixed-sulf = unidentified sulfate phases

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Table 19: Closed-system secondary mineralogy as seen in VNIR, XRD, and SEM-EDS VNIR XRD SEM-EDS Large platy Fe-oxide with elongated hexahydrite and anhydrite crystals, Hexahydrite, Epsomite, Hexahydrite, Epsomite, Hematite, Solids with a coating of amorphous mixed Fe-sulfate, Na/Al-sulfate, Maghemite, Goethite/Hematite Maghemite, Anhydrite Mg/Al-sulfate, Ti-oxide, and SiO2 Large masses of hexahydrite mixed with amorphous SiO2, some small

SMBxl Hexahydrite, Epsomite, Dried Gel Hexahydrite, Goethite areas with Fe/Al/K-sulfate (Fe-alunite), Na/Al-sulfate (natroalunite), and Maghemite, Goethite needles of Fe-oxide Substrate of basaltic glass with a coating of crystalline natrojarosite with Hexahydrite 84% Fe Natrojarosite, Alunogen, Solids varying Fe levels, sheets of Alunogen and hexahydrite, platy anhydrite, 65 – 92% Fe Natrojarosite Magnesiochloritoid, Epsomite small spherical Fe-oxide

SMBgl Titanomagnetite, Pickeringite, Quartz, Large platy hexahydrite, and some amorphous mixed with SiO2; platy Dried Gel Pickeringite, Alunogen 84% Fe Natrojarosite, FeSi pickeringite and alunogen, and unidentified Fe-sulfate (rhomboclase?) Jarosite, Cubic and dipyramidal natroalunite with varying Fe substitution growing Anhydrite, Hexahydrite, Epsomite, Solids 25% Fe Natroalunite, on large platy anhydrite, with amorphous coatings of Jarosite, Alunite

Hexahydrite SiO2 + Mg/Al-sulfates BAS - Hexahydrite, Alunogen, 25% Fe Natroalunite, Abundant cubic and dipyramidal natroalunite with varying

HN Dried Gel 25 – 49% Fe Natroalunite, Mg/Al-sulfate Fe substitution, anhydrite, amorphous coating of Mg/Al-sulfate + SiO2 Anhydrite

25 – 49% Fe Natroalunite, Alunogen, Anhydrite, Cubic Fe-bearing natroalunite and platy anhydrite, with Solids

Mg/Al-sulfate 25% Fe Natroalunite amorphous SiO2 + Mg/Al/Ca-sulfate coating

BAS - 25 – 49% Fe Natroalunite, 25% Fe Natroalunite, Platy anhydrite crystals with amorphous Dried Gel CN Mg/Al-sulfate Alunogen, Anhydrite Mg/Al-sulfate + SiO2 coating

Large platy anhydrite with abundant cubic & dipyramidal natroalunite 25 – 49% Fe Natroalunite, Anhydrite, Alunogen, Solids crystals (occasional slight Fe substitution), some amorphous Mg/Al-sulfate 20% Fe Natroalunite

Mg/Al-sulfate + SiO2 coating AND - 25 – 49% Fe Natroalunite, Hexahydrite, Alunogen, Platy anhydrite with cubic natroalunite and platy alunogen and B Dried Gel Mg/Al-sulfate 25% Fe Natroalunite hexahydrite, with amorphous Mg/Al-sulfate + SiO2 coating

Mg-sulfate, Solids Jarosite Amorphous mixed Mg-sulfate + SiO2 + alunite with some Fe substitution Amorphous SiO2 (sinter) Hexahydrite, Abundant platy anhydrite and alunogen, with amorphous mixture of SiO2 AND Dried Gel Anatase, Alunite, Jarosite Amorphous SiO2 (sinter) + hexahydrite + alunite (with and without significant Fe-substitution)

Substrate of SiO2 glass and quartz, with a few small areas of Solids Amorphous SiO2 (opal) No visible change K/Al-sulfate (alunite)

RHY Dried Gel Unidentified sulfate Unidentified sulfate Amorphous mixed Fe/K/Na/Al-sulfates + SiO2

Solids No visible change No visible change Small spots of amorphous SiO2 on surface of unaltered glass Analcime, Ferrinatrite, Pieces of unaltered glass, with small anhydrite and mixed OBS Dried Gel Not enough material MgSiO3, Amorphous SiO2 Na/Al/Mg/Fe-sulfates

Table 20: Mineral phases with a calculated SI > -0.4 in closed-system alteration fluids

BAS

AND

SMBxl SMBgl HN CN B AND RHY OBS

SiO2

Amorphous SiO2 SiO2 (0.1171) (0.0696) (0.1194) (0.0858) (0.1371) (0.1178) (0.1494) (0.0534) Chalcedony SiO2 1.1316 1.0841 1.1339 1.1003 1.1516 1.1323 1.1639 1.0679 Quartz SiO2 1.4028 1.3553 1.4051 1.3715 1.4228 1.4035 1.4351 1.3391 Cristobalite SiO2 0.8523 0.8048 0.8546 0.8210 0.8723 0.8530 0.8846 0.7886 Tridymite SiO2 1.2370 1.1895 1.2393 1.2057 1.2570 1.2377 1.2693 1.1733 Sulfates

Anhydrite Ca(SO4) (-0.2985) (-0.3235) (-0.2839) (-0.3348) (-0.1944)

Gypsum Ca(SO4) • 2 H2O (-0.1204) (-0.1454) (-0.1058) (-0.3210) (-0.1567) (-0.0163) (-0.3634) Alunite KAl3(SO4)2(OH)6 (0.2658) 1.2482

(Natro)Jarosite (Na/K)Fe3(SO4)2(OH)6 1.5939 2.4523 Oxides and Hydroxides

Gibbsite Al(OH)3 (-0.0684)

Todorokite (Na,Ca,K)2Mn6O12 • 3-4.5 H2O 12.213 9.4802

Pyrolusite MnO2 0.9145 0.6408 4.6199 4.2295 1.4106

Birnessite Mn2O4 • 1.5 H2O 14.7791 11.6558

Bixbyite (Mn,Fe)2O3 (-0.3295)

Hematite Fe2O3 (0.0934) 9.811 10.4297 0.9531

Goethite FeO(OH) 4.4272 4.7366 (-0.0017) Phyllosilicates and Zeolites

Beidellite Na0.5Al2.5Si3.5O10(OH)2 • n H2O 1.9224 2.6154

Nontronite Na0.3Fe2(Si,Al)4O10(OH)2 • n H2O 1.2916 1.9341 14.3698 14.9721 3.8177

Kaolinite Al2Si2O5(OH)4 2.4587 3.0989

Pyrophyllite Al2Si4O10(OH)2 3.6396 4.2125

Clinoptilolite (Na,K,Ca)2-3Al3(Al,Si)2Si13O36 • 12H2O 0.9693 1.1975 Phases at or near equilibrium: -0.4 < SI < 0.4; Denoted with “( SI )” Supersaturated phases: SI > 0.4

4.3.1.2. Glassy synthetic Mars basalt (SMBgl)

SMBgl alteration products (Figure 30d-f, Table 19) contained hexahydrite/epsomite and

anhydrite, as well as Al-sulfates (alunogen, pickeringite), Fe-oxides (titanomagnetite, and another

unidentified Fe-oxide), an unidentified Fe-silicate and Fe-sulfate, amorphous SiO2, and abundant

natrojarosite with ~84% Fe in the B site (alunite group: AB3(SO4)2(OH)6, with K/Na/Ca occupying

the A site, and Al/Fe in the B site) as identified in bulk powder XRD, using McCollom et al.’s

(2014) solid solution series. Fluids collected from the SMBgl vessel contained 1.50x106 ppb Fe –

an order of magnitude higher than all other samples (Figure 31), and are at equilibrium or

supersaturated with respect to hematite (SI = 0.0934) and nontronite (SI = 1.9341) (Table 20). Fe-

oxides/hydroxides were identified in the secondary mineral assemblage via XRD and SEM-EDS,

but as with the SMBxl alteration products, nontronite does not appear to be present, and may be

kinetically inhibited.

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4.3.1.3. High Fe terrestrial basalt (HN-BAS)

Alteration of the high Fe terrestrial basalt (HN-BAS) created products which appear similar

to the SMBxl and SMBgl secondary assemblages (Figure 30g, Table 19). SEM imaging of the

HN-BAS products shows a pervasive crust of natroalunite with varying Fe substitution coating

large platy anhydrite crystals, and patches of mixed Mg-sulfate + amorphous SiO2 coating.

However, the HN-BAS products lack any Fe-oxides/hydroxides, which are abundant in SMBxl

products. Additionally, intermediate alunite group minerals that are produced from HN-BAS

alteration contain less Fe in the B site compared to their SMBgl counterparts (average 25 % Fe in

HN-BAS products vs 84% Fe in SMBgl products). Despite the lack of Fe-oxides and hydroxides

in the alteration products, hematite, goethite, and the Mn/Fe-oxide bixbyite, are at equilibrium or

supersaturated with respect to the alteration fluid (Table 20), and are therefore expected to be

present in our secondary mineral assemblage. In addition to the supersaturated state of hematite

and goethite, nontronite is again supersaturated with respect to the fluid (Table 20), but lacking in

the secondary mineral assemblage, once again indicating possible kinetic barriers to formation.

4.3.1.4. Low Fe terrestrial basalt (CN-BAS)

CN-BAS products are dominated by Fe/Mg/Al-sulfates and amorphous SiO2 (Table 19).

In SEM images (Figure 30h, i), sulfates (with the exception of Fe-bearing natroalunite) are present

in a mixed microcrystalline coating along with SiO2. The microcrystalline nature is inferred based

on the presence of distinct peaks in the XRD patterns for hexahydrite, alunogen, and anhydrite –

indicating their presence in crystalline form rather than as mixed amorphous sulfates. As with the

other basaltic parent materials, CN-BAS alteration fluids are supersaturated with respect to

hematite, goethite, and nontronite (Table 20), despite the lack of these phases in the alteration

products.

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4.3.1.5. Terrestrial basaltic andesite (B-AND)

Alteration products from the close-system B-AND experiment consist primarily of ~20-

25% Fe natroalunite, hexahydrite, alunogen, anhydrite, and amorphous SiO2 (Table 19). SEM

imagery of these samples show large platy anhydrite crystals and dipyramidal natroalunite

encrusting an SiO2 substrate (Figure 30j). Although Fe-bearing natroalunite is abundant in the B-

AND alteration products, the overall level of Fe substitution is lower than those observed in SMBgl

and HN-BAS alunite group minerals. As with the basalt alteration fluids, hematite and nontronite

are also supersaturated in the B-AND fluid (Table 20), and goethite is at equilibrium; however, no

Fe-oxides/hydroxides or nontronite were detected in VNIR, XRD, or SEM-EDS (Table 19).

4.3.1.6. Terrestrial andesite (AND)

XRD analysis of the AND solid alteration products resulted in a diffraction pattern that

was nearly identical to the parent rock, with the addition of small jarosite peaks. Despite minimal

change in the XRD pattern, VNIR and SEM-EDS analysis show amorphous SiO2 mixed with Mg-

sulfate/hexahydrite, and Fe-bearing alunite in addition to the unaltered primary materials (Figure

30k, Table 19). Dried gel materials contain additional anatase, alunogen, and anhydrite (Figure

30l, Table 19). Unlike experiments with the more mafic parent materials (SMB, HN-BAS, CN-

BAS, and B-AND), mineral saturation states in the AND fluid are in agreement with the observed

mineral assemblages (Table 20), with amorphous SiO2 and anhydrite/gypsum at equilibrium with

the fluid. It is important to note, however, that the nearly unchanged XRD pattern for the AND

solid product indicates that this alteration reaction has not progressed very far, and is likely not

representative of an equilibrium state.

4.3.1.7. Terrestrial rhyolite (RHY) and obsidian (OBS)

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The most felsic endmembers of our experiments, RHY and OBS, displayed minimal

alteration of the parent material. XRD analysis of the RHY and OBS solid alteration products, and

VNIR of the OBS solid products showed no change from the parent rock diffraction pattern.

However, minor areas of amorphous SiO2 were observed coating nearly unaltered RHY and OBS

substrates through SEM-EDS analysis (Figure 30m – o, Table 19). Analysis of the RHY dried gel

materials identified Fe/K/Na/Al-sulfates (likely alunite group) mixed with amorphous SiO2, and

OBS dried gel was identified as anhydrite and mixed amorphous Mg/Fe/Na/Al-sulfates. No

additional minerals are saturated with respect to the fluids (Table 20).

4.3.2. Flow through system

4.3.2.1. Crystalline synthetic Mars basalt (SMBxl)

Simulated “flow through” alteration of our SMBxl parent material showed an evolution

from initial secondary mineral assemblages dominated by Fe-sulfates (e.g. rhomboclase, voltaite,

magnesioferrite) to a strongly leached SiO2/Ti/Al-dominated assemblage (Figure 32, Table 21).

By week 10, both XRD and SEM-EDS show a solid phase of primarily amorphous SiO2, with

small amounts of Fe-sulfates remaining in the dried gel. The Fe-sulfates in the dried gel persist

through the experiment with relative abundance (~17 – 18 wt % in the dried gel materials). VNIR

reflectance spectra of the final (week 20) SMBxl products also indicate the presence of Fe-sulfates,

with absorptions attributed to amorphous SiO2 (sinter) and rhomboclase.

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Figure 32: Flow through SMBxl alteration products in SEM; Si – amorphous SiO2; An – anhydrite; Pk – pickeringite; Hx – hexahydrite; Nat-j – natrojarosite; Nat-alun - natroalunite

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Table 21: Flow through SMBxl secondary mineralogy as seen in XRD and SEM-EDS XRD SEM-EDS Large Fe-sulfate substrate, with hexagonal platy Na/Fe-sulfate, natrojarosite, anhydrite, and hexahydrite Solids Hexahydrite, natrojarosite

with slight Fe substitution, dipyramidal and acicular Fe-sulfates, with a coating of amorphous SiO2 2 Quartz, Hexahydrite, Pickeringite, Diamond-shaped Fe- and Fe/Mg-sulfate (rhomboclase with occasional Mg substitution?), Week Week Dried Gel Alunite, Rhomboclase, Voltaite acicular pickeringite, anhydrite, and amorphous Mg/Fe-sulfate + SiO2 Substrate of platy anhydrite with Al/Mg-sulfates (hexahydrite, alunogen, pickeringite), Solids Hexahydrite, Alunogen, Anhydrite

natroalunite and natrojarosite, Fe-sulfate, and amorphous SiO2

4 Voltaite, Alunite, Tamarugite, Platy anhydrite and acicular Mg/Al-sulfate with diamond-shaped Fe/Mg-sulfate (Mg-copiapite Week Week Dried Gel Grunerite, Magnesioferrite, or Mg-rhomboclase?) and amorphous coatings of mixed SiO2 + Na/Al/Mg/Fe sulfates Cristobalite, Halotrichite Anhydrite, Cristobalite, Platy and acicular anhydrite with Fe/Ti-oxide (titanomagnetite?), Solids

Amorphous SiO2 Na/Al-silicate, Mg/Al-sulfates, with amorphous SiO2 coating + occasional Ti

6 Cristobalite, Anatase, Voltaite,

Week Week Dried Gel Magnesioferrite, Rhomboclase, Platy Fe-sulfate (rhomboclase?) with amorphous SiO2 + Al-sulfate Anhydrite

Solids Not enough material for XRD Platy anhydrite, Mg/Al-sulfates and Al-sulfates with amorphous mixed SiO2 + Mg/Al-sulfates

8 Rhomboclase, Abundant cubic Mg/Fe/Al-sulfates with varying Na/Ca (rhomboclase or magnesiovoltaite Week Week Dried Gel Amorphous SiO2 with varied substitution?), with amorphous SiO2 + Mg/Fe/Al/Na/Ca-sulfates Anatase, Cristobalite, Solids Amorphous SiO2 with occasional mixed Na/Ca/Al-sulfates

Amorphous SiO2

Cristobalite, SiO2, Large platy anhydrite, cubic Mg/Fe/Al-sulfates with varying Na/Ca (rhomboclase or magnesiovoltaite 10

Week Week Dried Gel Voltaite, Gmelinite-Ca, with varied substitution?), small Ti-rich area (anatase?) and possible gmelinite, with an Amorphous SiO2 amorphous mixed SiO2 +Mg/Na/Al-sulfate coating Anatase, Cristobalite, Solids Platy anhydrite with abundant amorphous SiO2 Amorphous SiO2

12 Massive amorphous SiO2 with mixed Ca/Na/Al-sulfate and occasional Ti, Week Week Dried Gel Anatase, Amorphous SiO2 platy anhydrite crystals scattered throughout Anatase, Cristobalite, 2

Solids Amorphous SiO , with a bit of anhydrite Amorphous SiO2 14 Platy anhydrite interspersed with massive coating of SiO2 + Mg/Al-sulfate, Week Week Dried Gel Not enough material for XRD small cubic crystal with Na/Al/Si (natrolite?)

Solids Not enough material for XRD Amorphous SiO , with occasional mixed Al-sulfate

16 SiO2, Anhydrite, Voltaite,

Week Week Dried Gel Platy anhydrite with amorphous Na/Al/Ca/Mg/Si coating Cristobalite, Amorphous SiO2

Solids Not enough material for XRD Amorphous SiO2, with occasional mixed Al-sulfate

Week Week Dried Gel Not enough material for XRD Amorphous SiO2 + Mg/Al-sulfate, occasionally with Na/Ca/Ti Cristobalite, Anatase, 2

Solids Amorphous SiO , with occasional mixed Al-sulfate Amorphous SiO2 20 Anhydrite, Rhomboclase, Diamond-shaped Fe-sulfate, with amorphous SiO2 + Na/Al/Ca/Mg-sulfate, Week Week Dried Gel Cristobalite, Amorphous SiO2 and occasional Ti (anatase?)

The evolution of the SMBxl fluid chemistry (Figure 33) shows a rapid influx of most cations (with the exception of Ti and K, which are located almost exclusively in accessory oxides and glass) as the major primary mineral phases are leached of their mobile cations. Si(aq) concentrations remain consistently at ~105 ppb throughout the experiment due to the rapid

5 precipitation of amorphous SiO2 when [Si(푎푞)] reaches 10 ppb (Iler, 1979; Jonckbloedt, 1998).

Any additional release of Si from the primary materials results in further SiO2(am) precipitation on their surfaces. Mg and Fe are immediately released to the week 2 alteration fluid as the olivine is rapidly dissolved, indicated by the [Mg] and [Fe] that are one to several orders of magnitude higher than all other cations in week 2. Rapid olivine dissolution is confirmed via SEM-EDS and XRD

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of week 2 alteration products – no olivine is detected by either method. Continued plagioclase dissolution is visible in the increase in [Na] and [Al] between weeks 2 and 4. Despite the increased addition of Ca to the fluid from plagioclase and pyroxene dissolution, [Ca] does not increase due to the precipitation of abundant crystalline anhydrite, which is at equilibrium with the alteration fluid (Table 22). [Ti] and [K] do not increase significantly until week 4, due to their sequestration in the accessory Fe/Ti-oxides and basaltic glass. Their notable lack in week 2 fluids indicate delayed dissolution of the accessory oxides and glass. Despite its predicted thermodynamic instability, previous experiments have shown basaltic glass to be largely unreactive (Marcucci and

Hynek, 2014; McCollom et al., 2013b), and our delayed glass dissolution is in agreement with this previous work. Saturation index calculations show supersaturation of hematite, goethite, and nontronite (Table 22); however, none of these phases were identified in our alteration assemblages.

As in the closed-system experiments, this may be due to kinetic inhibitors in the case of nontronite, and partitioning of Fe into the low-solubility Fe-bearing alunite group minerals (Baron and Palmer,

1996) rather than Fe-oxide/hydroxide.

Figure 33: Major element fluid chemistry evolution in flow through experiments; Lack of data points for [K+] in SMBxl are due to concentrations below instrument detection limits.

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Table 22: Mineral phases with a calculated SI > -0.4 in flow through SMBxl alteration fluids

2 Week 4 Week 6 Week 8 Week 10 Week 12 Week 14 Week 16 Week 18 Week 20 Week SiO2 Amorphous SiO2 SiO2 (0.1295) (0.0759) (0.0806) (0.0106) (-0.0366) (-0.0003) (0.0060) (0.0018) (0.0840) (0.0326) Chalcedony SiO2 1.1440 1.0904 1.0951 1.0251 0.9779 1.0142 1.0205 1.0163 1.0985 1.0471 Quartz SiO2 1.4152 1.3616 1.3663 1.2963 1.2491 1.2854 1.2917 1.2875 1.3697 1.3183 Cristobalite SiO2 0.8647 0.8111 0.8158 0.7458 0.6986 0.7349 0.7412 0.7370 0.8192 0.7678 Tridymite SiO2 1.2494 1.1958 1.2005 1.1305 1.0833 1.1196 1.1259 1.1217 1.2039 1.1525 Sulfates Anhydrite Ca(SO4) (-0.1748) (-0.1302) (-0.3778) (-0.3842) (-0.3993) Gypsum Ca(SO4) • 2 H2O (0.0033) (0.0479) (-0.1997) (-0.2061) (-0.2806) (-0.2745) (-0.2711) (-0.2212) Oxides and Hydroxides Pyrolusite MnO2 0.7301 Hematite Fe2O3 1.9922 Goethite FeO(OH) 0.5179 Phyllosilicates Nontronite Na0.3Fe2(Si,Al)4O10(OH)2 • n H2O 3.9369 0.8401 Phases at or near equilibrium: -0.4 < SI < 0.4; Denoted with “( SI )” Supersaturated phases: SI > 0.4

4.3.2.2. Glassy synthetic Mars basalt (SMBgl)

Flow through alteration of the SMBgl resulted in an initial secondary mineral assemblage

consisting of Fe- and Mg-sulfates (hexahydrite, pickeringite, voltaite, rhomboclase), anhydrite,

and amorphous SiO2 (Figure 34, Table 23). After 4 weeks of alteration, solid alteration products

transitioned to primarily Al- and Ca-sulfates, anatase, and amorphous SiO2. Fe-sulfates were no

longer present in the solid products, but persisted in the dried gel throughout the duration of the

experiment. Final (week 20) dried gel products contained highly crystalline Fe-sulfates which were

not present in the XRD pattern, indicating their presence at < 1 wt %, below the XRD detection

limit. VNIR analysis – which is capable of identifying trace components at < 1 wt % (Black and

Hynek, 2017; Marcucci et al., 2013) – of the week 20 materials produced a reflectance spectra

containing absorptions indicative of Fe-sulfates voltaite and/or rozenite.

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Figure 34: Flow through SMBgl Alteration products in SEM; Gl – unaltered glass; Si – amorphous SiO2; An – anhydrite; Pk – pickeringite; Hx – hexahydrite

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Table 23: Flow through SMBgl secondary mineralogy as seen in XRD and SEM-EDS XRD SEM-EDS Anhydrite, Hexahydrite, Platy anhydrite, rod-like hexahydrite, Fe/Mg-sulfate ridges, and a pickeringite crust, with a Solids

Amorphous SiO2 mixed amorphous SiO2 + Na/Al/Mg-sulfate coating with some Ti 2 Voltaite, Rhomboclase, Acicular pickeringite with cubic and/or dipyramidal Mg/Fe/Al-sulfate and diamond-shaped Week Week Dried Gel Na-Alum, Pickeringite Mg/Fe-sulfate (Mg-rhomboclase?), with an amorphous SiO2 coating

Solids Anhydrite, Amorphous SiO2 Strongly etched SiO2 glass substrate with coatings of Al-sulfate, anhydrite, and hexahydrite

4 Platy anhydrite, with Al/Mg-sulfate needles (pickeringite?), dipyramidal Anhydrite, Voltaite, Week Dried Gel Mg/Fe/Al-sulfate, small spherical Fe-sulfate, platy Al-sulfate, with mixed amorphous Tamarugite, Amorphous SiO2 SiO2 + Mg/Fe/Al/Na/Ca-sulfates, with occasional Ti enrichment

Solids Anhydrite, Amorphous SiO2 Platy anhydrite with grains of etched SiO2 glass

6 Rhomboclase, Voltaite, Platy anhydrite with cubic/dipyramidal Mg/Fe/Al/Ca-sulfate, radiating Week Week Dried Gel Anhydrite, Amorphous SiO2 platy Fe-sulfate, and amorphous SiO2 coating Al-sulfate with small crystals of anhydrite, massive hexahydrite, and anatase coatings, Solids Anhydrite, Amorphous SiO2

ridges of Na/Al-sulfates (natroalunite?) filling cracks in SiO2 glass substrate 8

Week Week Dried Gel Not enough material for XRD Large platy anhydrite with cubic Mg/Al/Fe-sulfates (Mg-rhomboclase?) and amorphous SiO2

Solids Amorphous SiO2 SiO2 glass substrate with small areas of Al/Ca-sulfate, and Ti-oxide

Week Week Dried Gel Not enough material for XRD Platy anhydrite and cubic Mg/Fe/Al/Ca-sulfates (Mg-rhomboclase?) with amorphous SiO2

Solids Amorphous SiO2 Strongly etched SiO2 glass with minor Al-sulfate + amorphous SiO2

12 Platy anhydrite, hexagonal plates of Fe-sulfate, cubic Mg/Fe/Al-sulfate, and Week Week Dried Gel Not enough material for XRD Al-sulfate with amorphous coatings of SiO2 + Mg-sulfate + Ti-oxide

Solids Amorphous SiO2 Strongly etched SiO2 glass, minor mixed Al-sulfates + SiO2 and Al-silicate

14 Platy anhydrite, with flaky coating of Ti-oxide, spherical and acicular Fe-sulfates, Week Week Dried Gel Not enough material for XRD and an amorphous coating of mixed SiO2 + Ti-oxide + Al-sulfate

Solids Amorphous SiO2 Strongly etched SiO2 glass

Week Week Dried Gel Not enough material for XRD Platy anhydrite with Ti-oxide (some areas with slight Fe) with amorphous SiO2 + Al-sulfate coating

Solids Amorphous SiO2 SiO2 glass with a hint of Al-sulfate

Week Week Dried Gel Not enough material for XRD Hexagonal plates of Fe-sulfate with amorphous mixture of SiO2 + Ti-oxide + Al-sulfate

Solids Amorphous SiO2 SiO2 glass

Week Week Dried Gel Anatase, Amorphous SiO2 Platy diamond and hexagonal-shaped Fe-sulfate with amorphous SiO2 + Ti-oxide coating

Fluid chemistry in the SMBgl reaction is indicative of a rapid dissolution of olivine by week 2, resulting in high [Mg] and [Fe] in the week 2 fluid. Remaining cations (Al, Ca, K, Mn, and Na) follow a relatively parallel linear trend as they are quickly leached from the basaltic glass in the initial stages. The glass becomes depleted as the experiment progresses, contributing fewer cations to the alteration fluid. [Mg] decreases more rapidly than the other cations due to the precipitation of abundant Mg-sulfates (hexahydrite and pickeringite). Similar to the SMBxl fluid,

[Ca] and [Si] remain at relatively constant levels for the majority (for Ca) or entirety (for Si) of the experiment. Saturation indices (Table 24) for amorphous SiO2 and anhydrite/gypsum show these phases to be at equilibrium with respect to the fluid. [Ca] and [Si] remain relatively constant at

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these equilibrium levels due to the precipitation of anhydrite and SiO2, which is confirmed by

XRD and SEM-EDS (Table 23). The continued increase of [Ti] from week 2 to week 4 may be due to rapid neutralization of [H+] in the first two weeks of the experiment, as significant amounts of cations are leached from the glass, quickly increasing the pH from ~0 to pH 1. This cation release may increase the pH enough to prevent further Ti mobility, as Ti has been shown to be a mobile element in very limited environments below pH 3 (e.g. Furnes, 1984; Minitti et al., 2007).

Additional Ti is leached from the glass in the subsequent weeks as fewer cations are available to neutralize the acidity of the alteration fluid.

Table 24: Mineral phases with a calculated SI > -0.4 in flow through SMBgl alteration fluids

2 Week 4 Week 6 Week 8 Week 10 Week 12 Week 14 Week 16 Week 18 Week 20 Week SiO2 Amorphous SiO2 SiO2 (0.1095) (0.0388) (0.0792) (0.0381) (-0.0438) (0.0810) (0.0493) (0.0492) (0.0455) (0.0500) Chalcedony SiO2 1.1240 1.0533 1.0937 1.0526 0.9707 1.0955 1.0638 1.0637 1.0600 1.0645 Quartz SiO2 1.3952 1.3245 1.3649 1.3238 1.2419 1.3667 1.3350 1.3349 1.3312 1.3357 Cristobalite SiO2 0.8447 0.7740 0.8144 0.7733 0.6914 0.8162 0.7845 0.7844 0.7807 0.7852 Tridymite SiO2 1.2294 1.1587 1.1991 1.1580 1.0761 1.2009 1.1692 1.1691 1.1654 1.1699 Sulfates Anhydrite Ca(SO4) (-0.2960) (-0.2307) (-0.3184) (-0.3165) (-0.3100) Gypsum Ca(SO4) • 2 H2O (-0.1179) (-0.0526) (-0.1403) (-0.1384) (-0.1323) Phyllosilicates Nontronite Na0.3Fe2(Si,Al)4O10(OH)2 • n H2O 0.7197 0.2231 Phases at or near equilibrium: -0.4 < SI < 0.4; Denoted with “( SI )” Supersaturated phases: SI > 0.4

4.4. Discussion

4.4.1. Effect of parent lithology on secondary mineralogy

Results from our closed-system reactions show variations in secondary mineral assemblages as parent FeOT increases (Table 25), even within the primary basalts alone. Secondary mineralogy resulting from alteration of our 19.8 wt % FeOT crystalline “Mars” basalt was dominated by Fe-oxides and hydroxides, with additional Fe present in a mixed amorphous sulfate

+ SiO2 coating and Fe-bearing alunite. Despite the abundance of Fe-oxides/hydroxides in the secondary SMB assemblages, and the documentation of these phases in terrestrial analog sites

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(Black and Hynek, 2017; e.g. Hynek et al., 2013; Marcucci et al., 2013), no Fe-oxides/hydroxides were identified in any of our altered terrestrial materials.

Table 25: Summary of secondary mineral phases in all experimental setups. Bold font indicates Fe-bearing secondary minerals Flow through system Closed system

SMBxl SMBgl

BAS

AND

SMBxl SMBgl HN CN B AND RHY OBS 2 Week 4 Week 6 Week 8 Week 10 Week 12 Week 14 Week 16 Week 18 Week 20 Week 2 Week Week 6 Week 8 Week 10 Week 12 Week 14 Week 16 Week 18 Week 20 Week SiO2 & silicates Amorphous SiO2 X X X X X X X X X X X X X X X X X X X X X X X X X X X X Quartz - X ------X ------Cristobalite ------X - X X X X - X ------FeSi - X ------Grunerite ------X ------MgSiO3 ------X ------Magnesiochloritoid - X ------Sulfates Anhydrite X X X X X X - - X X X X X X X X - X X X X X X X X X - - Tamarugite ------X ------X ------Na-Alum ------X ------Hexahydrite/Epsomite X X X - X X - - X X ------X X - X ------Pickeringite - X ------X - - X ------X X ------Alunogen - X X X X X - - - X - X ------X - X - X - - X - (Natro)alunite X - X - X X X - X X ------X ------Fe-alunite (undefined %) X - X X - X ------20% Fe natroalunite - - - - X ------25% Fe natroalunite - - X X X ------84% Fe natroalunite - X ------(Natro)jarosite - X X - - X - - X X ------Rhomboclase - ? ------X X X X ? - - - - X X - X X X X X - X X Voltaite ------X X X ? X - - X - - X X X ------Ferrinatrite ------X ------Mixed amorphous sulfates (Fe-bearing) X - - - - - X - X X - X ------X ------Mixed amorphous sulfates (non-Fe-bearing) - - X X X X - - - - X X X X X X X X X - - - X X X X - - Oxides & Hydroxides Anatase X - - - - X - - - - X - X X X - ? X - - - X X X X X X X Goethite X ------Hematite X ------Maghemite X ------Magnesioferrite ------X X ------Titanomagnetite - X ------X ------Unidentified Fe-oxide/hydroxide - X ------Zeolites Gmelinite-Ca ------X ------Analcime ------X ------

The additional Fe in the SMBxl and SMBgl parent rocks results in Fe-enriched primary olivine and pyroxene (Figure 28), the dissolution of which contribute an increased proportion of

Fe to the alteration fluid that becomes available for secondary mineral precipitation.

+ 2+ 2+ (푀푔, 퐹푒)2푆푖푂4 + 4퐻 → 푛 푀푔 + 푛 퐹푒 + 푆푖(푂퐻)4 Equation 5: Dissolution of olivine in acid

+ 2+ + 3+ 3+ 3+ 4+ (퐶푎, 푁푎)(푀푔, 퐹푒, 퐴푙, 푇푖)(푆푖, 퐴푙)2푂6 + 6퐻 → 푛 퐶푎 + 푛 푁푎 + 푛 푀푔 + 푛 퐹푒 + 푛 퐴푙 + 푛푇푖 + 푛 푆푖(푂퐻)4 Equation 6: Dissolution of pyroxene in acid

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Average olivine compositions measured by EPMA-WDS show elevated levels of Fe substitution in the SMBxl parent materials relative to our SMBgl, HN-BAS, and CN-BAS basalts.

+ 2+ 2+ SMBxl average olivine: (푀푔1.24퐹푒0.76)푆푖푂4 + 4퐻 → 1.24 푀푔 + 0.76 퐹푒 + 푆푖(푂퐻)4 + 2+ 2+ SMBgl average olivine: (푀푔1.56퐹푒0.37)푆푖푂4 + 4퐻 → 1.56 푀푔 + 0.37 퐹푒 + 푆푖(푂퐻)4 + 2+ 2+ HN-BAS average olivine: (푀푔1.58퐹푒0.40)푆푖푂4 + 4퐻 → 1.58 푀푔 + 0.40 퐹푒 + 푆푖(푂퐻)4 + 2+ 2+ CN-BAS average olivine: (푀푔1.39퐹푒0.59)푆푖푂4 + 4퐻 → 1.39 푀푔 + 0.59 퐹푒 + 푆푖(푂퐻)4

Similar to olivine, EPMA-WDS analysis of our primary basalts shows in increased level of Fe substitution in SMBxl pyroxenes relative to HN-BAS and CN-BAS (no pyroxenes are present in

SMBgl materials). From our observed pyroxene compositions, 0.71 mol Fe is released for every mol of SMBxl dissolution compared to 0.27 and 0.29 mol Fe for HN-BAS and CN-BAS dissolution, respectively. Therefore, the complete dissolution of the Fe-enriched SMBxl olivine and pyroxene results in an overall 67 - 119% increase in mol Fe released to the alteration fluid relative to our terrestrial basalts (HN-BAS and CN-BAS). This additional influx of Fe ions to the solution likely creates a supersaturated state for the common secondary Fe-bearing minerals, jarosite (or intermediate alunite group minerals), hematite, and goethite, all of which are present in our SMBxl alteration products (Table 19,Table 21, Table 25). Despite the supersaturated state of hematite and goethite in the HN-BAS and CN-BAS fluids (Table 20), they are notably absent from the alteration products. This may be due to the low solubility of jarosite (Baron and Palmer,

1996), which should result in preferential rapid precipitation of the Fe-sulfate over Fe- oxides/hydroxides. Indeed, jarosite or an intermediate alunite group mineral is present in each of our basalt-hosted experiments. The lack of hematite and goethite in the HN-BAS and CN-BAS assemblages may be due to the partitioning of Fe into sulfates, preventing the direct precipitation of Fe-oxides/hydroxides. It is possible that the additional Fe(aq) in the SMBxl fluid resulted in increased jarosite precipitation, which was then transformed to the observed hematite and goethite

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(Barrón et al., 2006). Or, alternatively, the increased Fe(aq) allowed for precipitation of both jarosite-type sulfates and Fe-oxides/hydroxides.

4.4.2. Evolution of flow through SMB alteration

Flow through alteration of our SMBxl and SMBgl resulted in initial Fe-dominated secondary mineral assemblages, consisting of both Fe-bearing alunite group minerals and additional Fe-sulfates, rhomboclase and voltaite. Despite the initial abundance of Fe-sulfates, flow through alteration assemblages quickly transitioned to amorphous SiO2, Al-sulfate, and Ti-oxide dominated. Rhomboclase and voltaite were still present in the dried gel materials at ~17 – 18 wt

% throughout the duration of the experiment. However, no Fe-sulfates were detected in the solids.

The presence of these lingering Fe-sulfates in the gel material may be due to the continued to slow leaching of Fe from basaltic glass as the reaction progressed. [Si] in the flow through fluids remained at the saturation state of 105 ppb (Iler, 1979; Jonckbloedt, 1998) throughout the experiment, with amorphous SiO2 rapidly precipitating on the surface of the parent materials as soon as it was leached. Any Fe being leached from the parent materials at this time may have become incorporated into the amorphous gel coating (which was later separated and dried) rather than new crystalline material. The evolution of secondary mineral assemblages suggests that similar flow through hydrothermal settings on Mars may result in alteration assemblages that are not enriched in Fe-bearing secondary mineralogy, despite the original high-Fe parent basalt, due to rapid leaching and removal of Fe from the system. Hydrothermally altered Martian deposits which are enriched in Fe-bearing secondary minerals may indicate areas that were altered over short timeframes, or operated as nearly closed systems (e.g. localized hot springs or fumaroles).

Major element cation concentrations from the alteration fluids taken at each 2-week interval show two distinct phases of alteration (Figure 35). Fluid compositions during the initial phase –

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from experiment initiation to week 6 – are dominated by the dissolution of primary minerals. After week 6, primary minerals are notably absent from XRD patterns, with the exception of plagioclase feldspar in the SMBxl materials. After week 6, fluid composition evolves as a function of glass dissolution, with rates of increase for cations roughly equal to observed molar ratios in the primary glass (e.g. phase 2 SMBgl Ti/Fe rate = 0.040, SMBgl primary glass Ti/Fe = 0.047). Notable exceptions to this are Mg/Fe and Na/Al, which follow a linear trend for the duration of the experiment. Na/Al increases at a similar rate (slope m) in the SMBxl (0.352) and SMBgl (0.365) fluids, which is roughly equal to the observed Na/Al molar ratio in the SMBgl glass (0.272) and

SMBxl glass (0.463, calculated) and plagioclase (0.322, observed) – indicating that Na and Al concentrations in the SMBgl alteration fluid are controlled predominantly by glass dissolution, while Na and Al concentrations in the SMBxl fluid are controlled by both glass and plagioclase dissolution. The slightly elevated Na/Al for the SMBgl fluids may indicate a preferential leaching of Na from the primary glass, or more likely, the ratio may be artificially skewed higher due to rapid precipitation of Al in the form of Al-sulfates such as alunogen and alunite group minerals.

Mg/Fe concentrations increase in a similar linear fashion, with the rate of increase for Mg/Fe in

SMBxl fluids at 0.852 and SMBgl at 1.050. The SMBgl rate is roughly equal to the observed

Mg/Fe ratios in the SMBgl glass (1.055), which is likely the primary control on Mg/Fe concentration in the SMBgl fluid, as the parent material is ~75 volume % glass. Mg/Fe ratios within the SMBgl olivines are significantly higher (4.216), but the olivines make up a much smaller portion of the primary material (Table 17), and therefore exert less influence on the fluid chemistry. Despite the clear linear trend for Mg/Fe in the SMBxl fluid (m = 0.852), these rates do not correlate with any of our observed SMBxl primary mineral or glass Mg/Fe ratios, indicating either a retention of Mg or Fe in primary materials, or a rapid removal of the leached ions from the

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fluid as they precipitate as secondary minerals such as Fe/Mg-sulfates and Fe-oxides/hydroxides.

Tosca et al. (2004) observed a retention of Mg in their primary pyroxene; however, our Mg/Fe

SMBxl pyroxene ratio is 0.648. A retention of Mg should result in a rate (m) lower than the observed Mg/Fe ratio, which does not agree with our observed fluid composition. The SMBxl olivine Mg/Fe ratio is 1.632, which could result in the observed rate (m) if Mg was preferentially retained. However, olivine is thermodynamically unstable at our experimental pH and temperature

(Jonckbloedt, 1998), and thus, dissolution should be congruent, though it may be slower than expected, as previous studies have shown it to be kinetically stable and present beyond plagioclase and pyroxene dissolution (Marcucci and Hynek, 2014). Therefore, we believe the SMBxl Mg/Fe ratio is affected by the removal of leached ions during secondary mineral precipitation.

Differences in the glassy or crystalline nature of the parent basalts appear to result in minimal differences in overall secondary mineral assemblages. Despite differences in mineral phases early in the SMBxl experiment (Table 21Table 25), late-stage secondary mineral assemblages appear the same, although there are variations in the relative abundance or crystalline nature of the minerals (Figure 32Figure 34). Initial variations in mineral assemblages may be attributed to initial dissolution of primary plagioclase and pyroxene in the SMBxl experiment contributing additional Ca, Na, Mg, Fe, and Al to the alteration fluid, which is confirmed in the fluid geochemistry. In the initial stages of alteration (weeks 0 – 6) Fe, Mg, Na are released to the alteration fluid at rates that exceed those in the SMBgl experiment (Figure 35). The additional Na and Fe result in the precipitation of natrojarosite, while the additional Mg precipitates as magnesioferrite along with the already abundant hexahydrite/epsomite (Table 25). The Ca-sulfate anhydrite is ubiquitous throughout both experiments, making it difficult to quantify an increase due to additional Ca leaching. The presence of hexahydrite in our solid alteration products for both the closed and flow-through systems (Table 19Table 21Table 23) is a notable contrast to the experiments by Marcucci and Hynek (2014), where the Mg-sulfates only formed as an evaporate product. While MgOT is higher in our SMB parent rocks than the 1999 Cerro Negro cinders used by Marcucci and Hynek (2014), hexahydrite is also found in our closed system CN-BAS solids

(Table 19), despite a similar MgOT abundance to Marcucci and Hynek’s (2014) parent material.

Our results indicate hexahydrite formation may not be limited to evaporate deposits in hydrothermal systems, and is capable of forming as an initial crystalline phase in both open and flow through systems.

The lack of Fe-oxides/hydroxides and the Fe-smectite nontronite despite their supersaturated state in several of the alteration fluids (Table 20Table 22Table 24) indicate kinetic

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inhibitors to their formation within our experimental setups. Goethite, hematite, and nontronite have all been identified in naturally-occurring terrestrial hydrothermal systems (Black and Hynek,

2017; e.g. Ehlmann et al., 2012; Hynek et al., 2013; Marcucci et al., 2013), and goethite and hematite were identified in our closed-system SMBxl alteration products. Peretyazhko et al. (2018) also produced dioctahedral smectite (e.g. montmorillonite, nontronite) in laboratory alteration of a synthetic glassy Adirondack-type basalt. XRD peaks indicate a montmorillonite composition rather than nontronite, although this could not be confirmed by VNIR or EPMA analyses. Despite the common presence of these minerals in alteration assemblages, and their supersaturated state in our alteration fluids, leached Fe in our alteration systems is partitioned into Fe-sulfates (alunite group, rhomboclase, and voltaite) or remains in solution.

4.4.3. Application to Martian hydrothermal systems

Fe-enriched basalt-hosted hydrothermal systems on Mars may produce similar mineralogical assemblages as are seen in our SMB alteration products. Initial alteration of Fe- enriched crystalline Martian basalts should rapidly release a large amount of Fe to the hydrothermal system from dissolution of primary olivine, pyroxene, and plagioclase. This increased Fe in the alteration fluid (relative to similar sites on Earth) may produce co-located Fe- oxides/hydroxides and Fe-sulfates, such as those forming in our closed and flow-through SMBxl systems. The lack of co-located Fe-oxides/hydroxides and Fe-sulfates in our SMBgl and terrestrial alteration products, as well as alteration deposits from Iceland (only 8% contain both Fe- oxides/hydroxides and Fe-sulfates) (Black et al., 2018) suggest the transition to mixed Fe- oxide/hydroxide/sulfate formation in hydrothermal acid-sulfate systems is between 17.0 – 19.8 wt

% FeOT in the parent basalt. Previous hydrothermal alteration experiments by Peretyazhko et al.

(2018) produced hematite and smectite from a glassy Adirondack composition (~19 wt% FeOT)

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basalt. Alteration of a 20.4 wt % FeOT synthetic glassy LA shergottite by Hurowitz et al. (2005) did produce Fe-sulfates and Fe-oxides, but not together. We attribute Hurowitz et al.’s (2005) lack of mixed Fe-oxides/sulfates despite the 20.4 wt % parent FeOT to the glassy nature of the parent material, which previous experiments have shown should react slowly (Marcucci and Hynek,

2014; McCollom et al., 2013b). Both closed and flow-through alteration of our 17.0 wt % FeOT

SMBgl produced only Fe-sulfates through the slower (relative to SMBxl alteration) release of Fe to the fluid from glass dissolution. Similarly slow dissolution of Hurowitz et al.’s (2005) glassy basalt may not allow [Fe] to become high enough in the alteration fluid to overcome competing mineral precipitation.

In situ investigation of the putative hydrothermal system at Home Plate in Gusev crater has resulted in identifications of Fe-sulfates, oxides, and hydroxides (Morris et al., 2008; Yen et al.,

2008). Compositions of unaltered basalts in the Gusev region range from 13.1 wt % FeOT at

Backstay (Ming et al., 2008) to 21.09 wt % FeOT at Adirondack (McSween et al., 2006). Fe-sulfate and Fe-oxide containing soils at Husband Hill (Morris et al., 2008) are similar to our SMBxl alteration products, and originate from comparatively similar parent basalt FeOT abundances. The nearest unaltered basalt to the Husband Hill Fe-sulfates and oxides is the Irvine basalt, which is also of similar composition to our SMBxl parent material (19.2 and 19.7 wt % FeOT, respectively)

(Ming et al., 2008). Observations of the nearby Paso Robles soils located further down the northern slope of the Columbia Hills identified 37 wt % Fe-sulfates with 5 wt % hematite (Yen et al., 2008).

However, the nearest unaltered basalt is the Backstay float rock sample, which only contained 13.1 wt % FeOT (Ming et al., 2008). From our experimental results, alteration of the Backstay basalt should result in an Fe-sulfate bearing assemblage which lacks Fe-oxides. Its presence only as a float rock (McSween et al., 2006) does not allow for the assumption that it is representative of the

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bedrock at this location, and secondary mineral assemblages indicate there may indeed be a more

Fe-rich primary basalt which extends further down the slopes of the Columbia Hills than the Irvine sample near the summit.

4.4.4. Implications for terrestrial analog studies

As expected, closed-system alteration of our SMB reveal a trend of increasing abundance of secondary Fe-bearing mineralogy as parent rock Fe increases. However, naturally-occurring field sites such as those in Iceland and Costa Rica (Black et al., 2018) rarely operate as true closed systems, making our closed system experiments an upper estimate for secondary Fe-bearing mineralogy abundance (Figure 36). Our experimental results suggest a negligible increase in overall abundance of secondary Fe-bearing mineralogy as basalts become increasingly enriched in primary Fe, although the mineralogy itself appears to vary (Fe-oxides/hydroxides vs. Fe-sulfates).

We therefore expect alteration of Fe-rich Martian basalts that occurs with limited loss of major element cations (e.g. fumaroles and isolated hot springs with limited fluid flow out of the system) to produce secondary assemblages in the range of ~20 – 32 wt % Fe-bearing minerals – similar to deposits that should form from alteration of terrestrial basalts and basaltic andesites.

4.5. Conclusions

We have investigated secondary mineral assemblages resulting from both closed and open- system acid-sulfate alteration using synthetic crystalline and glassy Martian basalts and terrestrial parent rocks spanning the mafic to felsic range. Closed-system secondary mineralogy of high Fe basalts is enriched in Fe-oxides/hydroxides and Fe-sulfates (alunite group, rhomboclase, and voltaite), while alteration of terrestrial parent materials only resulted in abundant alunite group sulfates, despite the supersaturation of Fe-oxides/hydroxides in the alteration fluid. Open-system alteration of our SMB parent materials began with abundant Fe-sulfates (alunite group, rhomboclase, and voltaite), but quickly transitioned to predominantly amorphous SiO2, Al-sulfate, and Ti-oxide as more soluble cations were leached out of the system at low pH. Continued exsolution of Fe from the basaltic glass resulted in persistent formation of Fe-sulfates rhomboclase and voltaite in the dried gel materials, possibly due to entrainment of the leached Fe cations in the saturated and rapidly precipitating amorphous SiO2. In all the basalt and basaltic andesite closed system experiments, as well as the early stages of the flow through SMB experiments, the Fe- smectite nontronite was supersaturated in the alteration fluid, but was not identified in any of the mineral assemblages. This is indicative of a kinetic barrier to nontronite formation in our experimental setups, despite its presence at several terrestrial analogs (Black and Hynek, 2017; e.g. Ehlmann et al., 2012; Hynek et al., 2013; Marcucci et al., 2013).

In both closed and flow-through systems, basalt alteration products were dominated by Fe- sulfates. Alteration fluids were supersaturated in Fe-oxides/hydroxides; however, these minerals only formed in SMBxl reactions. This suggests kinetic inhibitors to the formation of Fe- oxides/hydroxides such as competitive mineral precipitation favoring the formation of Fe-sulfates.

The lack of Fe-oxides/hydroxides in SMBgl reactions indicates that the required parent FeOT to

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overcome this barrier is between 17.0 and 19.8 wt % FeOT. Altered deposits in the Columbia Hills in Gusev crater contain both Fe-sulfates and Fe-oxides/hydroxides, despite the nearby float rocks containing 13.1 wt % FeOT (Backstay). We believe the combined Fe-sulfates and oxides/hydroxides near Backstay are due to a more Fe-rich parent rock at this location that is not represented by Backstay, which may be impact ejecta from a different location altogether. The

Irvine sample from the summit of Columbia Hills contains 19.7 wt % FeOT and may be more indicative of the bulk bedrock in the region. Bulk powder XRD of closed-system secondary mineralogy suggests an upper estimate of ~20 – 32 total wt % Fe-bearing mineralogy from alteration of high-Fe Mars-type basalts. This is similar to abundances that formed from closed- system alteration of our terrestrial basalts and basaltic andesite.

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5. Discussion and Conclusions

The presence of hydrothermal systems on early Mars is easily inferred from the number of volcanic centers and widespread evidence for crustal water in the Noachian and Hesperian. In our search for Martian biosignatures relic hydrothermal systems are a key area of investigation due to their teeming biota in terrestrial systems as well as a high potential for biosignature preservation through precipitation of opaline silica. In our search for biosignatures and potentially habitable environments on Mars, it is necessary to interpret the geochemical conditions at the time these hydrothermal systems were active. To do so, we must be able to accurately and thoroughly characterize the secondary mineralogy as well as understand the geochemical processes which form those minerals in Martian hydrothermal systems. In this dissertation work, I contribute to our understanding of Martian hydrothermal systems by investigating both instrumentation methods to determine a “best practices” for hydrothermal sites on Mars, and the role of parent rock Fe content on secondary mineral assemblages. This work helps to better constrain the geochemical processes which may occur in iron-rich basalt-hosted Martian hydrothermal systems and adds to our overall body of knowledge regarding the mineralogy of hydrothermal alteration deposits.

In chapter 2, I investigated the best payloads and instrumentation methods for examining relic hydrothermal deposits on Mars by using Mars-analog instruments: VNIR, XRD, and Raman

Laser Spectrometry. Each of these investigative methods is currently employed or is planned for in situ site characterization on Mars. VNIR reflectance spectroscopy has been a mineralogical workhorse of planetary exploration since the arrival of the Mars Express OMEGA instrument in

2003 and MRO’s CRISM in 2006. Since the dawn of orbital VNIR observations, VNIR has also been incorporated into in rover science payloads for in situ investigation. Also included is rover payloads has been XRD – a key to understanding bulk mineralogy of rover sites, and a commonly-

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utilized mineralogical identification method in terrestrial studies. Raman laser spectrometry is a new technique for Mars investigation, and has not yet been deployed as part of an active mission.

However, Raman spectrometers are planned for both the upcoming Mars 2020 and ExoMars rovers in an effort to bring new capabilities to rover investigations.

For each of these instruments, I utilized an analogous terrestrial method to characterize the mineralogy in 100 hydrothermally altered samples from Iceland, Costa Rica, and Nicaragua.

Samples contained a wide range of secondary mineralogy, including amorphous and crystalline

SiO2, elemental sulfur, Fe/Mg/Ca/Al-sulfates, Fe/Ti-oxides and hydroxides, phyllosilicates, and zeolites. Samples were collected from 2013 field campaigns and covered all visible color and morphological variations across a site, to increase the likelihood that all alteration minerals were present in our returned samples. This study was the first to systematically interpret a large set of hydrothermally altered samples with a variety of rover-deployed instruments, with the goal of developing an instrument “best practices” for current and future rover-based Mars missions. Key findings from this work include:

1) For more robust phyllosilicate detection, future rover payloads should carry

hyperspectral VNIR spectrometers which cover the 300 – 2500 nm range. This will

allow for “ground truthing” of orbital CRISM and OMEGA data and detection of

phyllosilicates beyond contact science.

2) Raman spectroscopy is extremely useful in detection of Fe-oxides/hydroxides,

elemental sulfur, anatase, and crystalline sulfates, but the Fe-rich Martian soils and rock

coatings provide a challenge as Fe-rich materials tend to burn when in contact with the

laser, and elemental sulfur overwhelms the raman signal.

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3) Phyllosilicates may be detected in XRD without any additional processing when they

are present at ≥ 10 wt %. Samples containing less abundant phyllosilicates may display

characteristic peaks for phyllosilicates, but these are often masked by highly crystalline

materials such as sulfates.

4) Amorphous silica and phyllosilicate-bearing deposits are easily missed or mis-

interpreted when observed with the limited rover VNIR (PanCam) bands. Several of

our SiO2 and/or phyllosilicate-bearing samples would not be identified as such when

deconvolved to MER/MSL PanCam bands as these wavelengths are dominated by the

spectra of Fe-bearing minerals such as hematite and goethite.

The second part of this dissertation research included the characterization of alteration mineralogy at hydrothermal sites in Iceland, chosen specifically due to their parent basalt compositions. Previous work has shown that the parent rock composition plays a significant role in the secondary mineral assemblage, and Icelandic basalts are among the closest terrestrial analogs for Martian basalts due to their high Fe content. The three sites chosen for this study were the Námafjall geothermal field, Þeistareykir geothermal field, and Nesjavallavir power plant in the

Hengill Volcanic Complex. We chose these sites for their parent rock composition – particularly

Fe abundance. Parent basalts ranged from 9.3 to 11.8 wt % FeOT and were being altered in both mudpot and fumarolic environments. In addition to investigating the role of parent rock composition on secondary mineral assemblage, this study was also the first to characterize the bulk mineralogy of surface alteration deposits at both Þeistareykir and two new locations in the Hengill

Volcanic Complex which are new potential Mars analog sites. Key findings from this work include:

166

1) Identifications of new mineral phases at Námafjall, including cristobalite, ~60 and 80

% Fe natroalunite, minamiite, voltaite, apjohnite, and heulandite.

2) Characterization of a new Mars analog site at Þeistareykir found differences between

high and low fluid:rock mudpots, indicating varied geochemical processes at these

sites, including a possible Fe enrichment at the low fluid:rock mudpot from a high

temperature region in the subsurface.

3) Alteration mineralogy at our Hengill-2 site showed a complex geochemical system

resulting from several coalescing streams with varied pH and temperature. Stream

appearance (color, texture) was strongly controlled by the secondary mineralogy, and

potentially biogeochemical processes, as indicated by the abundant biofilms and

streamers found throughout the site.

4) Identification of ~60 % Fe natroalunite, adding to the growing list of hydrothermal

systems containing naturally-occurring intermediate alunite group minerals. SEM-EDS

of this sample shows both > 10 µm and < 2 µm crystalline natroalunite with varying

Fe contents, indicating non-uniform Fe substitution.

5) Direct correlation between wt % parent FeOT and the average total wt % of Fe-bearing

secondary minerals. Projecting this trend out to Mars-composition basalts, we predict

10 – 25 wt % Fe-bearing secondary mineralogy to form from alteration of these high

Fe basalts.

For the third part of this work, I continued to investigate the effect of parent Fe on secondary mineralogy through the use of geochemical laboratory experiments. Synthetic glassy and crystalline Mars basalts with 17.0 and 19.8 wt % FeOT (respectively) were created at the

University of Alaska Fairbanks experimental petrology lab. These synthetic Mars basalts (SMB)

167

were then altered at 100ºC in 1M H2SO4 for up to 140 days. I also altered terrestrial parent rocks that ranged in composition from high Fe basalts (Hengill, Iceland), lower Fe basalts (Cerro Negro,

Nicaragua), and basaltic andesite (Turrialba, Costa Rica) to more evolved andesite (Flagstaff sill,

Colorado), rhyolite (Ruby Mt., Colorado) and obsidian (Millard County, Utah), making this the first study to systematically investigate the effect of parent rock composition and the first to alter both crystalline and glassy Mars composition basalts. Closed system experiments were run for 70 days using each of the parent rocks, while extended 140-day flow through experiments were also run using each of the synthetic Mars basalts. Key findings from these experiments include:

1) Secondary mineral assemblages for all basalts were dominated by Fe-sulfates;

however, only the crystalline synthetic Mars basalt also produced Fe-

oxides/hydroxides. This is suggestive of competitive mineral precipitation that may

favor Fe-sulfates over Fe-oxides/hydroxides and is overcome by additional Fe release

to the alteration fluid upon dissolution of the primary minerals in the crystalline SMB.

The lack of mixed Fe-sulfates and oxides/hydroxides in any of our experiments

indicates this transition is between 17.0 and 19.8 wt % FeOT.

2) During the first 6 weeks of flow through alteration, the fluid composition and

subsequent mineral precipitation are controlled by dissolution of the primary olivine,

pyroxene, and feldspar. This results in particularly Fe-rich fluids for the high Fe SMB

reactions, as the olivine and pyroxene release 67 – 119% more Fe upon dissolution

relative to the high Fe (11.8 wt % FeOT, Hengill) and low Fe (10.1 wt % FeOT, Cerro

Negro) terrestrial basalts.

3) Continued flow-through alteration of both crystalline and glassy SMB is controlled by

glass composition beyond week 6. At this point, secondary minerals transition to

168

Al/Ca/Mg-sulfates, SiO2, and TiO2, with lingering Fe-sulfates at lower abundances

than in the initial alteration stage. These alteration assemblages are similar to those that

are produced by alteration of lower Fe terrestrial basalts and basaltic andesites.

4) Alteration products from my closed system experiments may be taken as an upper

estimate for total wt % Fe-bearing secondary mineralogy. This trend, when projected

out to Mars-composition basalts results in an upper estimate of ~35 wt % Fe-bearing

secondary mineralogy, with our results from Icelandic field sites (which likely have

some removal of Fe as ground and surface water exits the system) setting a lower

estimate of ~15 wt % Fe-bearing secondary mineralogy.

5) Terrestrial analog sites may provide valuable insight into geochemical processes which

may have occurred on Mars. However, these sites have limitations when investigating

Fe-bearing mineralogy. Additional influx of Fe from dissolution of primary Martian

olivines and pyroxenes should result in increased abundances of Fe-bearing secondary

minerals, as well as more common instances of co-located Fe-sulfates and Fe-

oxides/hydroxides.

The work presented in this dissertation aims to improve our understanding of Martian hydrothermal systems through both improving instrumentation methods and an investigation of the role of parent rock Fe abundance on secondary Fe-bearing mineral assemblages. Martian hydrothermal systems are a key environment in our search for biosignatures, and it is necessary for us to understand the various geochemical processes which may occur in these systems. Through this dissertation research I have identified several weaknesses with current rover instrumentation methods, which should also aid in future rover payload and operations planning. In addition to instrumentation methods, I have also characterized the surficial alteration mineralogy at several

169

Mars analog hydrothermal sites in Iceland, including previously undescribed systems at the

Nesjavallavir power plant in the Hengill Volcanic Complex, and the Þeistareykir geothermal field.

This work was the start of an investigation into the role of parent basalt composition on secondary mineralogy, with a focus on Fe-bearing minerals. Laboratory-based experiments were also run to control variables more than is possible using terrestrial field sites. This combination of experimental and field investigation resulted in several predictions for alteration assemblages that would form from high Fe Martian basalts. The importance of this work is twofold. By systematically and experimentally altering parent rocks that contained a wide range of parent

FeOT, I was able to observe changes in alteration mineralogy which may be attributed to parent

Fe, and therefore will be important for our investigation of Martian hydrothermal systems. This work also serves as a validation of our analog studies using terrestrial hydrothermal systems, despite their lower FeOT in the parent basalts. Combined field and laboratory alteration of Icelandic basalts in addition to synthetic high Fe Martian basalts provides a valuable comparison for future investigations of Icelandic hydrothermal systems as Mars analogs.

The work included in this dissertation serves as a crucial step in our understanding of

Martian hydrothermal systems and their potential for habitability. Future investigations of Mars will rely on our ability to maximize scientific return through optimizing instrument payloads and through our understanding of geochemical processes on Earth and how those may be extrapolated to the high Fe basalts of Mars. This work provides new insight into both of these challenges through an investigation of instrument methods, a systematic study of parent Fe and secondary alteration mineralogy, and a validation of commonly-utilized terrestrial analog sites. The scientific contributions within this dissertation will help us move forward in our study of Mars’ geologic history and its potential for habitability.

170

References

Abramov, O., Kring, D.A., 2005. Impact-induced hydrothermal activity on early Mars. J. Geophys. Res. 110. doi:10.1029/2005JE002453

Adams, J.B., Smith, M.O., Johnson, P.E., 1986. Spectral mixture modeling: A new analysis of rock and soil types at the Viking Lander 1 Site. J. Geophys. Res. 91, 8098–8112. doi:10.1029/JB091iB08p08098

Al-Hanbali, H.S., Sowerby, S.J., Holm, N.G., 2001. Biogenicity of silicified microbes from a hydrothermal system: relevance to the search for evidence of life on earth and other planets. Earth Planet. Sci. Lett. 191, 213–218.

Allen, C.C., Albert, F.G., Chafetz, H.S., Combie, J., Graham, C.R., Kieft, T.L., Kivett, S.J., McKay, D.S., Steele, A., Taunton, A.E., Taylor, M.R., Thomas-Keprta, K.L., Westall, F., 2000. Microscopic physical biomarkers in carbonate hot springs: implications in the search for life on Mars. Icarus 147, 49–67. doi:10.1006/icar.2000.6435

Alpers, C.N., Rye, R.O., Nordstrom, D.K., White, L.D., King, B.-S., 1992. Chemical, crystallographic and stable isotopic properties of alunite and jarosite from acid-hypersaline Australian lakes. Chem. Geol. 96, 203–226.

Altheide, T.S., Chevrier, V.F., Noe Dobrea, E., 2010. Mineralogical characterization of acid weathered phyllosilicates with implications for secondary martian deposits. Geochim. Cosmochim. Acta 74, 6232–6248. doi:10.1016/j.gca.2010.08.005

Andrews-Hanna, J.C., Phillips, R.J., 2007. Hydrological modeling of outflow channels and chaos regions on Mars. J. Geophys. Res. 112. doi:10.1029/2006JE002881

Andrews-Hanna, J.C., Zuber, M.T., Arvidson, R.E., Wiseman, S.M., 2010. Early Mars hydrology: Meridiani playa deposits and the sedimentary record of Arabia Terra. J. Geophys. Res. 115. doi:10.1029/2009JE003485

Ann Hodges, C., Moore, H.J., 1979. The Subglacial Birth of Olympus Mons and Its Aureoles. J. Geophys. Res. 84, 8061–8074. doi:10.1029/JB084iB14p08061

Ármannsson, H., 2016. The fluid geochemistry of Icelandic high temperature geothermal areas. Appl. Geochemistry 66, 14–64. doi:10.1016/j.apgeochem.2015.10.008

Ármannsson, H., 2004. Chemical aspects of the exploration of the Theistareykir high-temperature geothermal area, N-E Iceland, in: Wanty & Seal II (Ed.), Water-Rock Interaction.

Ármannsson, H., Gíslason, G., Torfason, H., 1986. Surface exploration of the Theistareykir high- temperature geothermal area, Iceland, with special reference to the application of geochemical methods. Appl. Geochemistry 1, 47–64. doi:10.1016/0883-2927(86)90037-5

Ármannsson, H., Guðmundsson, Á., Steingrímsson, B.S., 1987. Exploration and development of the Krafla geothermal area. Jökull 37, 13–29.

171

Arvidson, R.E., Squyres, S.W., Bell, J.F., Catalano, J.G., Clark, B.C., Crumpler, L.S., de Souza, P.A., Fairén, A.G., Farrand, W.H., Fox, V.K., Gellert, R., Ghosh, A., Golombek, M.P., Grotzinger, J.P., Guinness, E. a, Herkenhoff, K.E., Jolliff, B.L., Knoll, A.H., Li, R., McLennan, S.M., Ming, D.W., Mittlefehldt, D.W., Moore, J.M., Morris, R. V, Murchie, S.L., Parker, T.J., Paulsen, G., Rice, J.W., Ruff, S.W., Smith, M.D., Wolff, M.J., 2014. Ancient aqueous environments at Endeavour crater, Mars. Science (80-. ). 343. doi:10.1126/science.1248097

ASD Inc., 2017. TerraSpec 4 User Manual: ASD Document 600691.

ASD Inc., 2015a. TerraSpec HALO Field Guide: ASD Document 600072.

ASD Inc., 2015b. TerraSpec HALO User Manual: ASD Document 600073.

Baker, L.L., Agenbroad, D.J., Wood, S.A., 2000. Experimental hydrothermal alteration of a martian analog basalt: Implications for martian meteorites. Meteorit. Planet. Sci. 35, 31–38. doi:10.1111/j.1945-5100.2000.tb01971.x

Baker, V.R., 1978. The Spokane Flood Controversy and the Martian Outflow Channels. Science (80-. ). 202, 1249–1256.

Baldridge, A.M., Calvin, W.M., 2004. Hydration state of the Martian coarse-grained hematite exposures: Implications for their origin and evolution. J. Geophys. Res. E Planets 109, 2–9. doi:10.1029/2003JE002066

Baldridge, A.M., Hook, S.J., Crowley, J.K., Marion, G.M., Kargel, J.S., Michalski, J.L., Thomson, B.J., de Souza Filho, C.R., Bridges, N.T., Brown, A.J., 2009. Contemporaneous deposition of phyllosilicates and sulfates: Using Australian acidic saline lake deposits to describe geochemical variability on Mars. Geophys. Res. Lett. 36. doi:10.1029/2009GL040069

Bandfield, J.L., 2002. Global Mineral Distributions on Mars 107. doi:10.1029/2001JE001510

Bandfield, J.L., Amador, E.S., Thomas, N.H., 2013. Extensive hydrated silica materials in western Hellas Basin, Mars. Icarus 226, 1489–1498. doi:10.1016/j.icarus.2013.08.005

Banin, A., Hah, F.X., Kan, I., Cieelsky, A., 1997. Acidic volatiles and the Mars Soil. J. Geophys. Res. 102, 13341–13356.

Barns, S.M., Fundyga, R.E., Jeffries, M.W., Pace, N.R., 1994. Remarkable archaeal diversity detected in a Yellowstone National Park hot spring environment. Proc. Natl. Acad. Sci. U. S. A. 91, 1609–1613.

Baron, D., Palmer, C.D., 1996. Solubility of jarosite at 4-35 ° C. Geochim. Cosmochim. Acta 60, 185–195.

Barrón, V., Torrent, J., Greenwood, J.P., 2006. Transformation of jarosite to hematite in simulated Martian brines. Earth Planet. Sci. Lett. 251, 380–385. doi:10.1016/j.epsl.2006.09.022

172

Barth-Wirsching, U., Holler, H., 1989. Experimental studies on zeolite formation conditions. Eur. J. Mineral. 1, 489–506.

Beckerman, L.G., 2016. Evaluating Mineralogy at Terrestrial Analogs for Early Mars: Detection and Characterization of Clays with XRD and Investigation of Iron Substitution in Natroalunite. University of Colorado Boulder.

Beegle, L., Bhartia, R., White, M., DeFlores, L., Abbey, W., Wu, Y.-H., Cameron, B., Moore, J., Fries, M., Burton, A., Edgett, K.S., Ravine, M.A., Hug, W., Reid, R., Nelson, T., Clegg, S., Wiens, R., Asher, S., Sobron, P., 2015. SHERLOC: Scanning Habitable Environments with Raman & Luminescence for Organics & Chemicals. IEEE Aerosp. Conf. Proc. doi:10.1109/AERO.2015.7119105

Beegle, L.W., Wilson, M.G., Abilleira, F., Jordan, J.F., Wilson, G.R., 2007. A Concept for NASA’s Mars 2016 Astrobiology Field Laboratory. Astrobiology 7, 545–577. doi:10.1089/ast.2007.0153

Bell, J.F., Malin, M.C., Caplinger, M.A., Ravine, M.A., Godber, A.S., Jungers, M.C., Rice, M.S., Anderson, R.B., 2012. Mastcam Multispectral Imaging on the Mars Science Laboratory Rover: Wavelength coverage and imaging strategies at the Gale Crater field site. Lunar Planet. Sci. Conf. 2541–2542.

Bell, J.F., Squyres, S.W., Herkenhoff, K.E., Maki, J.N., Arneson, H.M., Brown, D., Collins, S.A., Dingizian, A., Elliot, S.T., Hagerott, E.C., Hayes, A.G., Johnson, M.J., Johnson, J.R., Joseph, J., Kinch, K., Lemmon, M.T., Morris, R. V., Scherr, L., Schwochert, M., Shepard, M.K., Smith, G.H., Sohl-Dickstein, J.N., Sullivan, R.J., Sullivan, W.T., Wadsworth, M., 2003. Mars Exploration Rover Athena Panoramic Camera (Pancam) investigation. J. Geophys. Res. 108. doi:10.1029/2003JE002070

Bellucci, G., Altieri, F., Bibring, J.P., The OMEGA Team, 2004. The OMEGA Instrument on board Mars Express: First Results. Mem. della Soc. Astron. Ital. Suppl. 5, 27–32.

Benison, K.C., Bowen, B.B., Oboh-Ikuenobe, F.E., Jagniecki, E. a., LaClair, D. a., Story, S.L., Mormile, M.R., Hong, B.-Y., 2007. Sedimentology of Acid Saline Lakes in Southern Western Australia: Newly Described Processes and Products of an Extreme Environment. J. Sediment. Res. 77, 366–388. doi:10.2110/jsr.2007.038

Berger, G., Velde, B., 1992. Chemical parameters controlling the propylitic and argillic alteration process. Eur. J. Mineral. 4, 1439–1454.

Bibring, J.-P., Langevin, Y., Mustard, J.F., Poulet, F., Arvidson, R., Gendrin, A., Gondet, B., Mangold, N., Pinet, P., Forget, F., The OMEGA team, 2006. Global Mineralogical and Aqueous Mars History Derived from OMEGA/Mars Express Data. Science (80-. ). 312, 400– 404. doi:10.1126/science.1122659

Bishop, J.L., 2005. Hydrated Minerals on Mars. Water Mars Life 96, 65. doi:10.1007/978-3-540- 31538-4_4

173

Bishop, J.L., Murad, E., Lane, M.D., Mancinelli, R.L., 2004. Multiple techniques for mineral identification on Mars: A study of hydrothermal rocks as potential analogues for astrobiology sites on Mars. Icarus 169, 311–323. doi:10.1016/j.icarus.2003.12.025

Bishop, J.L., Noe Dobrea, E.Z., McKeown, N.K., Parente, M., Ehlmann, B.L., Michalski, J.R., Milliken, R.E., Poulet, F., Swayze, G.A., Mustard, J.F., Murchie, S.L., Bibring, J.-P., 2008. Phyllosilicate Diversity and Past Aqueous Activity Revealed at Mawrth Vallis, Mars. Science (80-. ). 321, 830–834.

Björke, J.K., 2010. Fluid-rhyolite interaction in geothermal systems, Torfajökull Iceland - secondary surface mineralogy and fluid chemistry upon phase segregation and fluid mixing. University of Iceland.

Black, S.R., Hynek, B.M., 2017. Characterization of terrestrial hydrothermal alteration products with Mars analog instrumentation: Implications for current and future rover investigations. Icarus. doi:10.1016/j.icarus.2017.10.032

Black, S.R., Hynek, B.M., McHenry, L.J., Glenister, C., Cameron, B.I., McCollom, T.M., Ludyan, J., 2018. Bulk mineralogy of surficial hydrothermal acid-sulfate deposits at Námafjall, Þeistareykir Geothermal Field, and Hengill Volcano, Iceland: Implications for the identification and interpretation of hydrothermal deposits on Mars, in: Lunar and Planetary Science Conference. pp. 1643–1644.

Blake, D., Vaniman, D., Achilles, C., Anderson, R., Bish, D., Bristow, T., Chen, C., Chipera, S., Crisp, J., Des Marais, D., Downs, R.T., Farmer, J., Feldman, S., Fonda, M., Gailhanou, M., Ma, H., Ming, D.W., Morris, R. V., Sarrazin, P., Stolper, E., Treiman, A., Yen, A., 2012. Characterization and calibration of the CheMin mineralogical instrument on Mars Science Laboratory. Space Sci. Rev. 170, 341–399. doi:10.1007/s11214-012-9905-1

Blake, D.F., Vaniman, D., Anderson, R., Bish, D., Chipera, S., Chemtob, S., Crisp, J., Des-Marais, D.J., Downs, R., Farmer, J., M. Gailhanou, Ming, D., Morris, D., Stolper, E., Sarrazin, P., Treiman, A., Yen, A., 2009. The CheMin mineralogical instrument on the Mars Science Laboratory mission, in: Lunar and Planetary Science Conference. pp. 1484–1485.

Blake, S., 1984. Magma mixing and hybridization processes at the alkalic, silicic, Torfajökull central volcano triggered by tholeiitic Veidivötn fissuring, south Iceland. J. Volcanol. Geotherm. Res. 22, 1–31. doi:10.1016/0377-0273(84)90033-7

Blank, C.E., Cady, S.L., Pace, N.R., 2002. Microbial Composition of Near-Boiling Silica- Depositing Thermal Springs throughout Yellowstone National Park. Appl. Environ. Microbiol. 68, 5123–5135. doi:10.1128/AEM.68.10.5123

Blasius, K.R., Cutts, J.A., Guest, J.E., Masursky, H., 1977. Geology of the Valles Marineris: First Analysis of Imaging From the Viking 1 Orbiter Primary Mission. J. Geophys. Res. 82, 4067– 4091.

Bodine, M.W., Fernalld, T.H., 1973. EDTA dissolution of gypsum, anhydrite, and Ca-Mg carbonates. J. Sediment. Petrol. 43, 1152–1156. doi:10.1306/74D72932-2B21-11D7-

174

8648000102C1865D

Bost, N., 2012. Geochemical and mineralogical analysis of Mars analogue materials and the creation of the International Space Analogue Rock Store (ISAR). Université d’Orléans.

Bost, N., Ramboz, C., LeBreton, N., Foucher, F., Lopez-Reyes, G., De Angelis, S., Josset, M., Venegas, G., Sanz-Arranz, a., Rull, F., Medina, J., Josset, J.-L., Souchon, a., Ammannito, E., De Sanctis, M.C., Di Iorio, T., Carli, C., Vago, J.L., Westall, F., 2015. Testing the ability of the ExoMars 2018 payload to document geological context and potential habitability on Mars. Planet. Space Sci. 108, 87–97. doi:10.1016/j.pss.2015.01.006

Boynton, W. V., Feldman, W.C., Mitrofanov, I.G., Evans, L.G., Reedy, R.C., Squyres, S.W., Starr, R., Trombka, J.I., D’Uston, C., Arnold, J.R., Englert, P.A.J., Metzger, A.E., Wanke, H., Bruckner, J., Drake, D.M., Shinohara, C., Fellows, C., Hamara, D.K., Harshman, K., Kerry, K., Turner, C., Ward, M., Barthe, H., Fuller, K.R., Storms, S.A., Thornton, G.W., Longmire, J.L., Litvak, M.L., Ton’Chev, A.K., 2004. The Mars Odyssey Gamma-Ray Spectrometer Instrument Suite. Space Sci. Rev. 110, 37–83.

Bramble, M.S., Mustard, J.F., Salvatore, M.R., 2017. The geological history of Northeast Syrtis Major, Mars. Icarus 293, 66–93. doi:10.1016/j.icarus.2017.03.030

Bridges, J.C., Henson, R.A., Vago, J.L., Loizeau, D., Williams, R.M.E., Hauber, E., Sefton-Nash, E., the ExoMars Landing Site Selection Working Group, 2016. ExoMars landing site characterization and selection, in: Lunar and Planetary Science Conference. pp. 2170–2171.

Bridges, J.C., Schwenzer, S.P., Leveille, R., Westall, F., Wiens, R.C., Mangold, N., Bristow, T., Edwards, P., Berger, G., 2015. Diagenesis and clay mineral formation at Gale Crater, Mars. J. Geophys. Res. E Planets 120, 1–19. doi:10.1002/2014JE004757

Bristow, T.F., Bish, D.L., Vaniman, D.T., Morris, R. V., Blake, D.F., Grotzinger, J.P., Rampe, E.B., Crisp, J.A., Achilles, C.N., Ming, D.W., Ehlmann, B.L., King, P.L., Bridges, J.C., Eigenbrode, J.L., Sumner, D.Y., Chipera, S.J., Moorokian, J.M., Treiman, A.H., Morrison, S.M., Downs, R.T., Farmer, J.D., Des Marais, D., Sarrazin, P., Floyd, M.M., Mischna, M.A., McAdam, A.C., 2015. The origin and implications of clay minerals from Yellowknife Bay, Gale crater, Mars. Am. Mineral. 100, 824–836. doi:10.2138/am-2015-5077CCBYNCND

Brophy, G.P., Scott, E.S., Snellgrove, R.A., 1962. Sulfate studies II. Solid solution between alunite and jarosite. Am. Mineral. 47, 112–126.

Brown, A.J., Sutter, B., Dunagan, S., 2008. The MARTE VNIR Imaging Spectrometer Experiment: Design and Analysis. Astrobiology 8, 1001–1011. doi:10.1089/ast.2007.0142

Brož, P., Hauber, E., 2012. A unique volcanic field in Tharsis, Mars: Pyroclastic cones as evidence for explosive eruptions. Icarus 218, 88–99. doi:10.1016/j.icarus.2011.11.030

Buckingham, W.F., Sommer, S.E., 1983. Mineralogical Characterization of Rock Surfaces Formed by Hydrothermal Alteration and Weathering - Application to Remote Sensing. Econ. Geol. 78, 664–674.

175

Bull, J.M., Minshull, T.A., Mitchell, N.C., Dix, J.K., Hardardottir, J., 2005. Magmatic and tectonic history of Iceland’s western rift zone at Lake Thingvallavatn. Bull. Geol. Soc. Am. 117, 1451–1465. doi:10.1130/B25729.1

Cabrol, N.A., Grin, E.A., Carr, M.H., Sutter, B., Moore, J.M., Farmer, J.D., Greeley, R., Kuzmin, R.O., Desmarais, D.J., Kramer, M.G., Newsom, H., Barber, C., Thorsos, I., Tanaka, K.L., Barlow, N.G., Fike, D.A., Urquhart, M.L., Grigsby, B., Grant, F.D., Goursac, O. De, 2003. Exploring Gusev Crater with Spirit: Review of science objectives and testable hypotheses. J. Geophys. Res. 108. doi:10.1029/2002JE002026

Carr, M.H., 2006. The Surface of Mars, Cambridge Planetary Science Series. doi:10.1086/518077

Carr, M.H., 1995. The Martian drainage system and the origin of valley networks and fretted channels. J. Geophys. Res. 100, 7479–7507.

Carr, M.H., 1979. Formation of Martian Flood Features by Release of Water From Confined Aquifers. J. Geophys. Res. 84, 2995–3007.

Carr, M.H., Head, J.W., 2010. Geologic history of Mars. Earth Planet. Sci. Lett. 294, 185–203. doi:10.1016/j.epsl.2009.06.042

Carson, G.L., 2015. Hydrothermal acid-sulfate alteration at Krafla and Namafjall, NE Iceland: Implications for Gusev Crater and Meridiani Planum, Mars 142.

Carter, J., Poulet, F., Bibring, J.-P., Mangold, N., Murchie, S., 2013. Hydrous minerals on Mars as seen by the CRISM and OMEGA imaging spectrometers: Updated global view. J. Geophys. Res. Planets 118, 831–858. doi:10.1029/2012JE004145

Catling, D.C., Moore, J.M., 2003. The nature of coarse-grained crystalline hematite and its implications for the early environment of Mars. Icarus 165, 277–300. doi:10.1016/S0019- 1035(03)00173-8

Cavanagh, P.D., Bish, D.L., Blake, D.F., Vaniman, D.T., Morris, R.V., Ming, D.W., Rampe, E.B., Achilles, C.N., Chipera, S.J., Treiman, A.H., Downs, R.T., Morrison, S.M., Fendrich, K.V., Yen, A.S., Grotzinger, J., Crisp, J.A., Bristow, T.F., Sarrazin, P.C., Farmer, J.D., Marais, D.J. Des, Stolper, E.M., Morookian, J.M., Wilson, M.A., Spanovich, N., Anderson, R.C., The MSL Science Team, 2015. Confidence Hills mineralogy and CheMin results from base of Mt. Sharp, Pahrump Hills, Gale crater, Mars, in: Lunar and Planetary Science Conference.

Chapman, M.G., Tanaka, K.L., 2001. Interior trough deposits on Mars: Subice volcanoes? J. Geophys. Res. 106, 10087–10100. doi:10.1029/2000JE001303

Chemtob, S.M., Rossman, G.R., 2014. Timescales and mechanisms of formation of amorphous silica coatings on fresh basalts at K ī lauea Volcano , Hawai ’ i. J. Volcanol. Geotherm. Res. 286, 41–54. doi:10.1016/j.jvolgeores.2014.08.029

Chevrier, V., Poulet, F., Bibring, J.-P., 2007. Early geochemical environment of Mars as determined from thermodynamics of phyllosilicates. Nature 448, 60–63.

176

doi:10.1038/nature05961

Christensen, P.R., Anderson, D.L., Chase, S.C., Clancy, R.T., Clark, R.N., Conrath, B.J., Kieffer, H.H., Kuzmin, R.O., Malin, M.C., Pearl, J.C., Roush, T.L., Smith, M.D., 1998. Results from the Mars Global Surveyor Thermal Emission Spectrometer. Science (80-. ). 279, 1692–1698.

Christensen, P.R., Bandfield, J.L., Hamilton, V.E., Ruff, S.W., Kieffer, H.H., Titus, T.N., Malin, M.C., Morris, R. V, Lane, M.D., Clark, R.L., Jakosky, B.M., Mellon, M.T., Pearl, J.C., Conrath, B.J., Smith, M.D., Clancy, R.T., Kuzmin, R.O., Roush, T., Mehall, G.L., Gorelick, N., Bender, K., Murray, K., Dason, S., Greene, E., Silverman, S., Greenfield, M., 2001. Mars Global Surveyor Thermal Emission Spectrometer experiment: Investigation description and surface science results. J. Geophys. Res. 106, 23823–23871.

Christensen, P.R., McSween, H.Y., Bandfield, J.L., Ruff, S.W., Rogers, a D., Hamilton, V.E., Gorelick, N., Wyatt, M.B., Jakosky, B.M., Kieffer, H.H., Malin, M.C., Moersch, J.E., 2005. Evidence for magmatic evolution and diversity on Mars from infrared observations. Nature 436, 504–509. doi:10.1038/nature04075

Christensen, P.R., Mehall, G.L., Silverman, S.H., Anwar, S., Cannon, G., Gorelick, N., Kheen, R., Tourville, T., Bates, D., Ferry, S., Fortuna, T., Jeffryes, J., O’Donnell, W., Peralta, R., Wolverton, T., Blaney, D., Denise, R., Rademacher, J., Morris, R. V., Squyres, S., 2003. Miniature Thermal Emission Spectrometer for the Mars Exploration Rovers. J. Geophys. Res. 108. doi:10.1029/2003JE002117

Christensen, P.R., Morris, R. V., Lane, M.D., Bandfield, J.L., Malin, M.C., 2001. Global mapping of Martian hematite mineral deposits: Remnants of water-driven processes on early Mars. J. Geophys. Res. 106, 23873. doi:10.1029/2000JE001415

Christensen, P.R., Wyatt, M.B., Glotch, T.D., Rogers, A.D., Anwar, S., Arvidson, R.E., Bandfield, J.L., Blaney, D.L., Budney, C., Calvin, W.M., Fallacaro, A., Fergason, R.L., Gorelick, N., Graff, T.G., Hamilton, V.E., Hayes, A.G., Johnson, J.R., Knudson, A.T., McSween, H.Y., Mehall, G.L., Mehall, L.K., Moersch, J.E., Morris, R. V., Smith, M.D., Squyres, S.W., Ruff, S.W., Wolff, M.J., 2004. Mineralogy at Meridiani Planum from the Mini-TES Experiment on the Opportunity Rover. Science (80-. ). 306, 1733–1739. doi:10.1126/science.1104909

Christiansen, E.H., Burt, D.M., Sheridan, M.F., Wilson, R.T., 1983. The Petrogenesis of Topaz Rhyolites from the Western United States. Contrib. to Mineral. Petrol. 83, 16–30.

Clark, B.C., Baird, A.K., Weldon, R.J., Tsusaki, D.M., Schnabel, L., Candelaria, M.P., 1982. Chemical Composition of Martian Fines. J. Geophys. Res. 87, 10059–10067.

Clark, B.R.N., Swayze, G.A., Wise, R.A., Livo, K.E., Hoefen, T.M., Raymond, F., Sutley, S.J., Geological, U.S., 2013. USGS Digital Spectral Library splib06a 1–24.

Clark, R.N., 1999. Spectroscopy of Rocks and Minerals, and Principles of Spectroscopy, 3rd ed, Remote Sensing for the Earth Sciences: Manual of Remote Sensing. John Wiley & Sons, Inc. doi:10.1111/j.1945-5100.2004.tb00079.x

177

Clark, R.N., King, T.V. V., Klejwa, M., Swayze, G.A., Vergo, N., 1990. High Spectral Resolution Reflectance Spectroscopy of Minerals. J. Geophys. Res. 95, 12653–12680. doi:10.1029/JB095iB08p12653

Clark, R.N., Swayze, G.A., Wise, R., Livo, E., Hoefen, T., Kokaly, R., Sutley, S.J., 2007. USGS digital spectral library splib06a.

Clegg, S.M., Wiens, R.C., Maurice, S., Gasnault, O., Forni, O., Sharma, S.K., Misra, A.K., Rull, F., Johnson, J.R., 2014. Remote Geochemical and Mineralogical Analysis with SuperCam on the Mars 2020 Rover and on Earth, in: American Geophysical Union, Fall Meeting. p. abstract #P11A-3753.

Clifford, S.M., 1993. A Model for the Hydrologic and Climatic Behavior of Water on Mars. J. Geophys. Res. 98, 10973–11016.

Coradini, A., Piccioni, G., Amici, S., Bianchi, R., Capaccioni, F., Capria, M.T., De Sanctis, M.C., Di Lellis, A.M., Espinasse, S., Federico, C., Fonti, S., Arnold, G., Atreya, S.K., Owen, T., Blecka, M., Bini, A., Cosi, M., Pieri, S., Tacconi, M., 2001. MA _ MISS: Mars Multispectral Imager for Subsurface Studies. Adv. Sp. Res. 28, 1203–2001.

Cousins, C.R., Griffiths, A.D., Crawford, I.A., Prosser, B.J., Storrie-Lombardi, M.C., Davis, L.E., Gunn, M., Coates, A.J., Jones, A.P., Ward, J.M., 2010. Astrobiological considerations for the selection of the geological filters on the ExoMars PanCam instrument. Astrobiology 10, 933– 951. doi:10.1089/ast.2010.0517

Cousins, C.R., Gunn, M., Prosser, B.J., Barnes, D.P., Crawford, I.A., Griffiths, A.D., Davis, L.E., Coates, A.J., 2012. Selecting the geology filter wavelengths for the ExoMars Panoramic Camera instrument. Planet. Space Sci. 71, 80–100. doi:10.1016/j.pss.2012.07.009

Craddock, R.A., Howard, A.D., 2002. The case for rainfall on a warm, wet early Mars. J. Geophys. Res. 107. doi:10.1029/2001JE001505

Crecraft, H.R., Nash, W.P., Evans, S.H., 1981. Late Cenozioc Volcanism at Twin Peaks, Utah: Geology and Petrology. J. Geophys. Res. 86, 10303–10320.

Crovisier, J.-L., Advocat, T., Dussossoy, J.-L., 2003. Nature and role of natural alteration gels formed on the surface of ancient volcanic glasses (Natural analogs of waste containment glasses). J. Nucl. Mater. 321, 91–109. doi:10.1016/S0022-3115(03)00206-X

Crumpler, L.S., Head, J.W., Aubele, J.C., 1996. Calderas on Mars: characteristics, structure, and associated flank deformation. Geol. Soc. London, Spec. Publ. 110, 307–348. doi:10.1144/GSL.SP.1996.110.01.24

Cull-Hearth, S., van Venrooy, A., Clark, M.C., Cvitkovic, A., 2016. Acid-sulfate mixtures from Río Tinto, Spain: Spectral masking relationships and implications for Mars. Icarus 271, 387– 399. doi:10.1016/j.icarus.2016.01.001

Des Marais, D.J., Nuth, J.A., Allamandola, L.J., Boss, A.P., Farmer, J.D., Hoehler, T.M., Jakosky,

178

B.M., Meadows, V.S., Pohorille, A., Runnegar, B., Spormann, A.M., Marais, D.E.S., Al, E.T., 2008. The NASA Astrobiology Roadmap. Astrobiology 8, 715–730. doi:10.1089/ast.2008.0819

Desborough, G.A., Smith, K.S., Lowers, H.A., Swayze, G.A., Hammarstrom, J.M., Diehl, S.F., Leinz, R.W., Driscoll, R.L., 2010. Mineralogical and chemical characteristics of some natural jarosites. Geochim. Cosmochim. Acta 74, 1041–1056. doi:10.1016/j.gca.2009.11.006

Di Achille, G., Hynek, B.M., 2010. Ancient ocean on Mars supported by global distribution of deltas and valleys. Nat. Geosci. 3, 459–463. doi:10.1038/ngeo891

Djokic, T., Van Kranendonk, M.J., Campbell, K.A., Walter, M.R., Ward, C.R., 2017. Earliest signs of life on land preserved in ca. 3.5 Ga hot spring deposits. Nat. Commun. 8. doi:10.1038/ncomms15263

Dreibus, G., Wänke, H., 1985. Mars, a volatile-rich planet. Meteoritics 20, 367–381.

Drever, J.I., 1997. The Geochemistry of Natural Waters. Prentice Hall, Inc., Upper Saddle River, N. J.

Economou, T., 2001. Chemical analyses of martian soil and rocks obtained by the Pathfinder Alpha Proton X-ray spectrometer. Radiat. Phys. Chem. 61, 191–197. doi:10.1016/S0969- 806X(01)00240-7

Edwards, C.S., Bandfield, J.L., Christensen, P.R., Fergason, R.L., 2009. Global distribution of bedrock exposures on Mars using THEMIS high-resolution thermal inertia. J. Geophys. Res. 114. doi:10.1029/2009JE003363

Ehlmann, B.L., Bish, D.L., Ruff, S.W., Mustard, J.F., 2012. Mineralogy and chemistry of altered Icelandic basalts: Application to clay mineral detection and understanding aqueous environments on Mars. J. Geophys. Res. 117. doi:10.1029/2012JE004156

Ehlmann, B.L., Edwards, C.S., 2014. Mineralogy of the Martian Surface. Annu. Rev. Earth Planet. Sci. 42, 291–315. doi:10.1146/annurev-earth-060313-055024

Ehlmann, B.L., Mustard, J.F., Clark, R.N., Swayze, G.A., Murchie, S.L., 2011a. Evidence for low- grade metamorphism, hydrothermal alteration, and diagenesis on Mars from phyllosilicate mineral assemblages. Clays Clay Miner. 59, 359–377. doi:10.1346/CCMN.2011.0590402

Ehlmann, B.L., Mustard, J.F., Murchie, S.L., Bibring, J.-P., Meunier, A., Fraeman, A. a., Langevin, Y., 2011b. Subsurface water and clay mineral formation during the early history of Mars. Nature 479, 53–60. doi:10.1038/nature10582

Ehlmann, B.L., Mustard, J.F., Swayze, G.A., Clark, R.N., Bishop, J.L., Poulet, F., Des Marais, D.J., Roach, L.H., Milliken, R.E., Wray, J.J., Barnouin-Jha, O., Murchie, S.L., 2009. Identification of hydrated silicate minerals on Mars using MRO-CRISM: Geologic context near Nili Fossae and implications for aqueous alteration. J. Geophys. Res. 114. doi:10.1029/2009JE003339

179

Ehlmann, B.L., Swayze, G.A., Milliken, R.E., Mustard, J.F., Clark, R.N., Murchie, S.L., Breit, G.N., Wray, J.J., Gondet, B., Poulet, F., Carter, J., Calvin, W.M., Benzel, W.M., Seelos, K.D., 2016. Discovery of alunite in Cross crater, Terra Sirenum, Mars: Evidence for acidic, sulfurous waters. Am. Mineral. 101, 1527–1542.

El-Maarry, M.R., Black, S.R., Hynek, B.M., McHenry, L.J., 2017. Mineralogy of Fumarolic Deposits from Iceland as Analogs for Ancient Hydrothermal Systems on Mars: Role of Temperature, in: Lunar and Planetary Science Conference. pp. 2870–2871. doi:10.1029/2009JE003548.

Ellison, E., 2017. Raman Microspectroscopy Lab [WWW Document]. Univ. Color. Boulder. URL http://spot.colorado.edu/~templeta/Templeton_Lab/Raman_Lab.html (accessed 7.20.01).

Eshaghpour, M., 2003. Borehole Geology and Alteration Mineralogy of Well HE-9 in Hellisheidi Geothermal Field, SW-Iceland.

Farrand, W.H., Bell, J.F., Johnson, J.R., Rice, M.S., Jolliff, B.L., Arvidson, R.E., 2014. Observations of rock spectral classes by the Opportunity rover’s Pancam on northern Cape York and on Matijevic Hill, Endeavour Crater, Mars. J. Geophys. Res. E Planets 119, 2349– 2369. doi:10.1002/2014JE004641

Farrand, W.H., Glotch, T.D., Rice, J.W., Hurowitz, J. a., Swayze, G. a., 2009. Discovery of jarosite within the Mawrth Vallis region of Mars: Implications for the geologic history of the region. Icarus 204, 478–488. doi:10.1016/j.icarus.2009.07.014

Fassett, C.I., Head, James, W., 2007. Valley formation on martian volcanoes in the Hesperian: Evidence for melting of summit snowpack, caldera lake formation, drainage and erosion on Ceraunius Tholus. Icarus 189, 118–135. doi:10.1016/j.icarus.2006.12.021

Fassett, C.I., Head, J.W., 2008. The timing of martian valley network activity Constraints from buffered crater counting. Icarus 195, 61–89. doi:10.1016/j.icarus.2007.12.009

Fassett, C.I., Head, J.W., 2006. Valleys on Hecates Tholus, Mars: origin by basal melting of summit snowpack. Planet. Space Sci. 54, 370–378. doi:10.1016/j.pss.2005.12.011

Fassett, C.I., Head, J.W., 2005. Fluvial sedimentary deposits on Mars: Ancient deltas in a crater lake in the Nili Fossae region. Geophys. Res. Lett. 32. doi:10.1029/2005GL023456

Foley, C.N., 2003. Final chemical results from the Mars Pathfinder alpha proton X-ray spectrometer. J. Geophys. Res. 108, 8096. doi:10.1029/2002JE002019

Foulger, G.R., Toomey, D.R., 1989. Structure and Evolution of the Hengill-Grensdalur Volcanic Complex, Iceland: Geology, Geophysics, and Seismic Tomography. J. Geophys. Res. 94, 17511–17522. doi:10.1029/JB094iB12p17511

Frey, H. V., 2003. Buried impact basins and the earliest history of Mars, in: Sixth International Conference on Mars. p. 3104.

180

Furnes, H., 1984. Chemical changes during progressive subaerial palagonitization of a subglacial olivine tholeiite hyaloclastite: A microprobe study. Chem. Geol. 43, 271–285.

Garvin, J.B., Sakimoto, S.E.H., Frawley, J.J., Schnetzler, C.C., Wright, H.M., 2000. Topographic Evidence for Geologically Recent Near-Polar Volcanism on Mars. Icarus 145, 648–652. doi:10.1006/icar.2000.6409

Gebrehiwot, K., Koestono, H., Franzson, H., Mortensen, A.K., 2010. Borehole Geology and Hydrothermal Alteration of Well HE-24 At Hellisheidi Geothermal Field,SW-Iceland, in: Proceedings World Geothermal Congress.

Gendrin, A., Mangold, N., Bibring, J.-P., Langevin, Y., Gondet, B., Poulet, F., Bonello, G., Quantin, C., Mustard, J., Arvidson, R., LeMouélic, S., 2005. Sulfates in Martian Layered Terrains: The OMEGA/Mars Express View. Science (80-. ). 307, 1587–1591. doi:10.1126/science.1109087

Geptner, A.R., Ivanovskaya, T.A., Pokrovskaya, E. V., 2005. Hydrothermal Fossilization of Microorganisms at the Earth’s Surface in Iceland. Lithol. Miner. Resour. 40, 505–520. doi:10.1007/s10987-005-0048-2

Geptner, A.R., Ivanovskaya, T.A., Pokrovskaya, E. V., Lyapunov, S.M., Savichev, A.T., Gorbunov, A. V., Gor’kova, N. V., 2007. Hydrothermally Altered Hyaloclastites at the Earth’s Surface in the Rift Zone of Iceland: Problem of the Biochemogenic Accumulation of Trace Elements. Lithol. Miner. Resour. 42, 453–476. doi:10.1134/S0024490207050045

Getaneh, E., 2001. Borehole Geology and Alteration Mineralogy in the Upper Half of Well HE-3, Hellisheidi, SW-Iceland.

Ghatan, G.J., Head, J.W., 2002. Candidate subglacial volcanoes in the south polar region of Mars: Morphology, morphometry, and eruption conditions. J. Geophys. Res. 107. doi:10.1029/2001JE001519

Glotch, T.D., Bandfield, J.L., Tornabene, L.L., Jensen, H.B., Seelos, F.P., 2010. Distribution and formation of chlorides and phyllosilicates in Terra Sirenum, Mars. Geophys. Res. Lett. 37, 1–5. doi:10.1029/2010GL044557

Golden, D.C., Ming, D.W., Morris, R. V., Mertzman, S.A., 2005. Laboratory-simulated acid- sulfate weathering of basaltic materials: Implications for formation of sulfates at Meridiani Planum and Gusev crater, Mars. J. Geophys. Res. 110. doi:10.1029/2005JE002451

Golden, D.C., Morris, R. V., Ming, D.W., Lauer, H. V., Yang, S.R., 1993. Mineralogy of three slightly palagonitized basaltic tephra samples from the summit of Mauna Kea, Hawaii. J. Geophys. Res. 98, 3401–3411. doi:10.1029/92JE02590

Golombek, M.P., Bridges, N.T., 2000. Erosion rates on Mars and implications for climate change: Constraints from the Pathfinder landing site. J. Geophys. Res. 105, 1841–1853.

Golombek, M.P., Grant, J.A., Crumpler, L.S., Greeley, R., Arvidson, R.E., Bell, J.F., Weitz, C.M.,

181

Sullivan, R., Christensen, P.R., Soderblom, L.A., Squyres, S.W., 2006. Erosion rates at the Mars Exploration Rover landing sites and long-term climate change on Mars. J. Geophys. Res. 111. doi:10.1029/2006JE002754

Golombek, M.P., Grant, J.A., Parker, T.J., Kass, D.M., Crisp, J.A., Squyres, S.W., Haldemann, A.F.C., Adler, M., Lee, W.J., Bridges, N.T., Arvidson, R.E., Carr, M.H., Kirk, R.L., Knocke, P.C., Roncoli, R.B., Weitz, C.M., Schofield, J.T., Zurek, R.W., Christensen, P.R., Fergason, R.L., Anderson, F.S., Rice, J.W., 2003. Selection of the Mars Exploration Rover landing sites. J. Geophys. Res. 108. doi:10.1029/2003JE002074

Gorevan, S.P., Myrick, T., Davis, K., Chau, J.J., Bartlett, P., Mukherjee, S., Anderson, R., Squyres, S.W., Arvidson, R.E., Madsen, M.B., Bertelsen, P., Goetz, W., Binau, C.S., Richter, L., 2003. Rock Abrasion Tool: Mars Exploration Rover mission. J. Geophys. Res. 108. doi:10.1029/2003JE002061

Goudge, T.A., Mustard, J.F., Head, J.W., Fassett, C.I., Wiseman, S.M., 2015. Assessing the mineralogy of the watershed and fan deposits of the Jezero crater paleolake system, Mars. J. Geophys. Res. Planets 120, 775–808. doi:10.1002/2014JE004782.Received

Greeley, R., Crown, D.A., 1990. Volcanic Geology of Tyrrhena Patera, Mars. J. Geophys. Res. 95, 7133–7149.

Greeley, R., Guest, J.E., 1987. Geologic Map of the Eastern Equatorial Region of Mars.

Greenberger, R.N., Mustard, J.F., Kumar, P.S., Dyar, M.D., Breves, E. a., Sklute, E.C., 2012. Low temperature aqueous alteration of basalt: Mineral assemblages of Deccan basalts and implications for Mars. J. Geophys. Res. 117. doi:10.1029/2012JE004127

Gregg, T.K.P., Farley, M.A., 2006. Mafic pyroclastic flows at Tyrrhena Patera, Mars: Constraints from observations and models. J. Volcanol. Geotherm. Res. 155, 81–89. doi:10.1016/j.jvolgeores.2006.02.008

Griffiths, A.D., Coates, A.J., Jaumann, R., Michaelis, H., Paar, G., Barnes, D., Josset, J.-L., The PanCam Team, 2006. Context for the ESA ExoMars rover: the Panoramic Camera (Pancam) instrument. Int. J. Astrobiol. 5, 269–275. doi:10.1017/S1473550406003387

Grindrod, P.M., West, M., Warner, N.H., Gupta, S., 2012. Formation of an Hesperian-aged sedimentary basin containing phyllosilicates in Coprates Catena, Mars. Icarus 218, 178–195. doi:10.1016/j.icarus.2011.11.027

Grove, T.L., 1981. Use of FePt alloys to eliminate the iron loss problem in 1 atmosphere gas mixing experiments: Theoretical and practical considerations. Contrib. to Mineral. Petrol. 78, 298–304. doi:10.1007/BF00398924

Gulick, V.C., 2001. Origin of the valley networks on Mars: a hydrological perspective. Geomorphology 37, 241–268.

Gulick, V.C., 1998. Magmatic intrusions and a hydrothermal origin for fluvial valleys on Mars. J.

182

Geophys. Res. 103, 19365–19387.

Hagerty, J.J., Newsom, H.E., 2003. Hydrothermal alteration at the Lonar Lake impact structure, India: Implications for impact cratering on Mars. Meteorit. Planet. Sci. 38, 365–381. doi:10.1111/j.1945-5100.2003.tb00272.x

Halevy, I., Head, J.W., 2014. Episodic warming of early Mars by punctuated volcanism. Nat. Geosci. 7, 865–868. doi:10.1038/NGEO2293

Hamilton, V.E., Hoehler, T., Eigenbrode, J., Rafkin, S., Withers, P., Ruff, S., Yingst, R.A., Lim, D., Whitley, R., Beaty, D., Diniega, S., Hays, L., Zurek, R., 2015. Mars Science Goals, Objectives, Investigations and Priorities: 2015 Mars Exploration Program Analysis Group (MEPAG) 75.

Hanna, J.C., Phillips, R.J., 2005. Hydrological modeling of the Martian crust with application to the pressurization of aquifers. J. Geophys. Res. 110. doi:10.1029/2004JE002330

Harris, J.K., Cousins, C.R., Gunn, M., Grindrod, P.M., Barnes, D., Crawford, I.A., Cross, R.E., Coates, A.J., 2015. Remote detection of past habitability at Mars-analogue hydrothermal alteration terrains using an ExoMars Panoramic Camera Emulator. Icarus 252, 284–300. doi:10.1016/j.icarus.2015.02.004

Harris, W., White, N., 2007. X-ray Diffraction Techniques for Soil Mineral Identification, in: Methods of Soil Analysis. Soil Science Society of America, Madison, WI.

Hartmann, W.K., Berman, D.C., 2000. Elysium Planitia lava flows: Crater count chronology and geological implications. J. Geophys. Res. 105, 15011–15025.

Hartmann, W.K., Neukum, G., 2001. Cratering Chronology and the Evolution of Mars. Space Sci. Rev. 96, 165–194.

Hauber, E., Platz, T., Reiss, D., Le Deit, L., Kleinhans, M.G., Marra, W.A., De Haas, T., Carbonneau, P., 2013. Asynchronous formation of Hesperian and Amazonian-aged deltas on Mars and implications for climate. J. Geophys. Res. E Planets 118, 1529–1544. doi:10.1002/jgre.20107

Hausrath, E.M., 2007. Basalt weathering on Earth and on Mars. Pennsylvania State University.

Hausrath, E.M., Golden, D.C., Morris, R. V., Agresti, D.G., Ming, D.W., 2013. Acid sulfate alteration of fluorapatite, basaltic glass and olivine by hydrothermal vapors and fluids: Implications for fumarolic activity and secondary phosphate phases in sulfate-rich Paso Robles soil at Gusev Crater, Mars. J. Geophys. Res. E Planets 118, 1–13. doi:10.1029/2012JE004246

Hausrath, E.M., Tschauner, O., 2013. Natural fumarolic alteration of fluorapatite, olivine, and basaltic glass, and implications for habitable environments on Mars. Astrobiology 13, 1049– 64. doi:10.1089/ast.2013.0985

183

Head, J., 2013. The Early Climate History of Mars: “Warm and Wet” or “Cold and Icy”?, in: EGU General Assembly 2013. p. EGU2013-10106-1.

Head, J.W., Marchant, D.R., 2003. Cold-based mountain glaciers on Mars: Western Arsia Mons. Geology 31, 641–644.

Head, J.W., Pratt, S., 2001. Extensive Hesperian-aged south polar ice sheet on Mars: Evidence for massive melting and retreat, and lateral flow and ponding of meltwater. J. Geophys. Res. 106, 12275–12299.

Heald, P., Foley, N.K., Hayba, D.O., 1987. Comparative Anatomy of Volcanic-Hosted Epithermal Deposits: Acid-Sulfate and Adularia-Sericite Types. Econ. Geol. 82, 1–26. doi:10.2113/gsecongeo.82.1.1

Helgadóttir, H.M., Franzson, H., Óskarsson, N., Grönvold, K., Steinþórsson, S., 2015. Hydrothermal Alteration of Pyroxene in the Hellisheiði Geothermal Field, SW-Iceland. Proc. World Geotherm. Congr.

Helgadóttir, H.M., Snaebjörnsdóttir, S.Ó., Níelsson, S., Gunnarsdóttir, S.H., Matthíasdóttir, T., Harðarson, B.S., Einarsson, G.M., Franzson, H., 2010. Geology and Hydrothermal Alteration in the Reservoir of the Hellisheiði High Temperature System, SW-Iceland. Proc. World Geotherm. Congr. doi:10.13140/2.1.1008.0648

Henley, R.W., Ellis, A.J., 1983. Geothermal Systems Ancient and Modern: A Geochemical Review. Earth Sci. Rev. 19, 1–50.

Herd, C.D.K., Papike, J.J., Brearley, A.J., 2001. Oxygen fugacity of martian basalts from electron microprobe oxygen and TEM-EELS analyses of Fe-Ti oxides. Am. Mineral. 86, 1015–1024.

Hernández, P. a., Pérez, N.M., Fridriksson, T., Egbert, J., Ilyinskaya, E., Thárhallsson, A., Ívarsson, G., Gíslason, G., Gunnarsson, I., Jónsson, B., Padrón, E., Melián, G., Mori, T., Notsu, K., 2012. Diffuse volcanic degassing and thermal energy release from Hengill volcanic system, Iceland. Bull. Volcanol. 74, 2435–2448. doi:10.1007/s00445-012-0673-2

Hill, K.S., Hansford, G.M., Vernon, D., Talboys, D.L., Ambrosi, R.M., Bridges, J.C., Hutchinson, I.B., 2011. The Mars-XRD Instrument for Exomars: Combined X-Ray Diffraction and Fluorescence Measurements, in: Lunar and Planetary Science Conference. pp. 2107–2108.

Hoblitt, R., Larson, E., 1975. Bearing on the Structural Evolution of the Northeastern Margin of the Front Range, Colorado. Geol. Soc. Am. Bull. 86, 237–242.

Hoke, M.R.T., Hynek, B.M., Di Achille, G., Hutton, E.W.H., 2014. The effects of sediment supply and concentrations on the formation timescale of martian deltas. Icarus 228, 1–12. doi:10.1016/j.icarus.2013.09.017

Hoke, M.R.T., Hynek, B.M., Tucker, G.E., 2011. Formation timescales of large Martian valley networks. Earth Planet. Sci. Lett. 312, 1–12. doi:10.1016/j.epsl.2011.09.053

184

Horgan, B., 2013. Evolved magma on Mars. Nat. Geosci. 6, 991–992. doi:10.1038/ngeo2010

Horgan, B.H.N., Smith, R.J., Cloutis, E.A., Mann, P., Christensen, P.R., 2017. Acid weathering of basalt and basaltic glass: I. Near-infrared spectra, thermal-infrared spectra, and implications for Mars. J. Geophys. Res. Planets. doi:10.1002/2016JE005111

Horneck, G., 2000. The microbial world and the case for Mars. Planet. Space Sci. 48, 1053–1063.

Hugenholtz, P., Pitulle, C., Hershberger, K.L., Pace, N.R., 1998. Novel Division Level Bacterial Diversity in a Yellowstone Hot Spring. J. Bacteriol. 180, 366–376.

Humayun, M., Nemchin, A., Zanda, B., Hewins, R.H., Grange, M., Kennedy, A., Lorand, J.-P., Göpel, C., Fieni, C., Pont, S., Deldicque, D., 2013. Origin and age of the earliest Martian crust from meteorite NWA 7533. Nature 503, 513–516. doi:10.1038/nature12764

Hunt, G.R., 1977. Spectral signatures of particulate minerals in the visible and near infrared, in: Geophysics. pp. 501–513. doi:10.1190/1.1440721

Hurowitz, J.A., McLennan, S.M., Lindsley, D.H., Schoonen, M.A.A., 2005. Experimental epithermal alteration of synthetic Los Angeles meteorite: Implications for the origin of Martian soils and identification of hydrothermal sites on Mars. J. Geophys. Res. 110. doi:10.1029/2004JE002391

Hurowitz, J.A., McLennan, S.M., Tosca, N.J., Arvidson, R.E., Michalski, J.R., Ming, D.W., Schroder, C., Squyres, S.W., 2006. In situ and experimental evidence for acidic weathering of rocks and soils on Mars. J. Geophys. Res. 111. doi:10.1029/2005JE002515

Hynek, B., 2016. The great climate paradox of ancient Mars. Geology 44, 879–880. doi:10.1130/focus102016.1

Hynek, B.M., Beach, M., Hoke, M.R.T., 2010. Updated global map of Martian valley networks and implications for climate and hydrologic processes. J. Geophys. Res. 115. doi:10.1029/2009JE003548

Hynek, B.M., McCollom, T.M., Marcucci, E.C., Brugman, K., Rogers, K.L., 2013. Assessment of environmental controls on acid-sulfate alteration at active volcanoes in Nicaragua: Applications to relic hydrothermal systems on Mars. J. Geophys. Res. Planets 118, 1–22. doi:10.1002/jgre.20140

Hynek, B.M., McCollom, T.M., McHenry, L.J., Alvarado, G.E., 2014. Assessing hydrothermal alteration on early Mars through analog environments in Nicaragua, Costa Rica, Iceland, and Hawaii, in: Lunar and Planetary Science Conference. pp. 2172–2173. doi:10.1038/ngeo990.

Hynek, B.M., Mccollom, T.M., Rogers, K.L., 2011. Cerro Negro volcano, Nicaragua: An assessment of geological and potential biological systems on early Mars. Geol. Soc. Am. Spec. Pap. 483, 279–285. doi:10.1130/2011.2483(18)

Hynek, B.M., Osterloo, M.K., Kierein-Young, K.S., 2015. Late-stage formation of Martian

185

chloride salts through ponding and evaporation. Geology 43, 787–790. doi:10.1130/G36895.1

Hynek, B.M., Phillips, R.J., Arvidson, R.E., 2003. Explosive volcanism in the Tharsis region: Global evidence in the Martian geologic record. J. Geophys. Res. 108. doi:10.1029/2003JE002062

Iler, R.K., 1979. The Chemistry of Silica. John Wiley & Sons, Inc., New York.

Israel, E.J., Arvidson, R.E., Wang, A., Pasteris, J.D., Jolliff, B.L., 1997. Laser Raman spectroscopy of varnished basalt and implications for in situ measurements of Martian rocks. J. Geophys. Res. 102, 28705–28716.

Jacobsen, M.J., Balic-Zunic, T., Mitolo, D., Katerinopoulou, A., Garavelli, A., Jakobsson, S.P., 2014. Oskarssonite, AlF3, a new fumarolic mineral from Eldfell volcano, Heimaey, Iceland. Mineral. Mag. 78, 215–222. doi:10.1180/minmag.2014.078.1.15

Jercinovic, M.J., Keil, K., Smith, M.R., Schmitt, R.A., 1990. Alteration of basaltic glasses from north-central British Columbia, Canada. Geochim. Cosmochim. Acta 54, 2679–2696.

Johnson, J.R., Bell, J.F., Bender, S., Blaney, D., Cloutis, E., DeFlores, L., Ehlmann, B., Gasnault, O., Gondet, B., Kinch, K., Lemmon, M., Le Mouélic, S., Maurice, S., Rice, M., Wiens, R.C., the MSL Science Team, 2015. ChemCam passive reflectance spectroscopy of surface materials at the Curiosity landing site, Mars. Icarus 249, 74–92. doi:10.1016/j.icarus.2014.02.028

Jonckbloedt, R.C.L., 1998. Olivine dissolution in sulphuric acid at elevated temperatures — implications for the olivine process, an alternative waste acid neutralizing process. J. Geochemical Explor. 62, 337–346.

Kahn, R., Lee, S., Martin, T.Z., Zurek, R.W., 1990. The Martian Dust Cycle. Bull. Am. Astron. Soc. 22, 1076.

Kallenbach, R., Geiss, J., Hartmann, W.K., 2001. Chronology and Evolution of Mars, 12th ed. Space Science Series of ISSI, Bern, Switzerland.

King, P.L., McLennan, S.M., 2010. Sulfur on mars. Elements 6, 107–112. doi:10.2113/gselements.6.2.107

Kirkland, L.E., Herr, K.C., Adams, P.M., 2003. Infrared stealthy surfaces: Why TES and THEMIS may miss some substantial mineral deposits on Mars and implications for remote sensing of planetary surfaces. J. Geophys. Res. 108. doi:10.1029/2003JE002105

Kite, E.S., Halevy, I., Kahre, M. a., Wolff, M.J., Manga, M., 2013. Seasonal melting and the formation of sedimentary rocks on Mars, with predictions for the Gale Crater mound. Icarus 223, 181–210. doi:10.1016/j.icarus.2012.11.034

Klein, H.P., 1978. The Viking Biological Experiments on Mars. Icarus 34, 666–674.

186

Klimchouk, A., 1996. The dissolution and conversion of gypsum and anhydrite. Int. J. Speleol. 25, 21–36.

Klingelhöfer, G., 2004. Mössbauer In Situ Studies of the Surface of Mars. Hyperfine Interact. 117– 124. doi:10.1007/s10751-005-9019-1

Klingelhöfer, G., Morris, R. V., Bernhardt, B., Schroder, C., Rodionov, D.S., Jr., P.A. de S., Yen, A., R. Gellert, 1, 6 E. N. Evlanov, 3 B. Zubkov, 3 J. Foh, 1, 7 U. Bonnes, 7 E. Kankeleit, 7 P. Gu, ¨tlich, 1, D. W. Ming, 2 F. Renz, 1 T. Wdowiak, 8 S. W. Squyres, 9 R. E. Arvidson, 2004. Jarosite and Hematite at Meridiani Planum from Opportunity’s Mossbauer Spectrometer. Science (80-. ). 306, 1740–1745. doi:10.1126/science.1104653

Komar, P.D., 1979. Comparisons of the Hydraulics of Water Flows in Martian Outflow Channels with Flows of Similar Scale on Earth. Icarus 37, 156–181.

Korablev, O., Ivanov, A., Mantsevich, S., Fedorova, A., Kuzmin, R.O., Evdokimova, N., Trokhimovsky, A., Basilevsky, A., Rodionov, D., Kiselev, A., Arefiev, V., Kozlova, T., Kozlov, O., Semena, N., Titov, A., Vyazovetskiy, N., Kalinnikov, Y., Bibring, J., Poulet, F., Esposito, F., Kilian, R., Flahaut, J., Cloutis, E., Schmitz, N., Griffiths, A., 2014. An Infrared Spectrometer at the Mast of ExoMars Rover, in: 40th COSPAR Scientific Assembly. doi:10.1017/S1473550406003387

Kristmannsdóttir, H., Tómasson, J., 1974. Nesjavellir: Hydrothermal alteration in a high temperature area, in: International Symposium on Water-Rock Interaction. Prague.

Lafuente, B., Downs, R.T., Yang, H., Stone, N., 2015. The power of databases: the RRUFF project, in: Armbruster, T., Danisi, R.M. (Eds.), Highlights in Mineralogical Crystallography. W. De Gruyter, Berlin, Germany, pp. 1–30. doi:10.1515/9783110417104-003

Lane, M.D., Morris, R. V, Mertzman, S. a, Christensen, P.R., 2002. Evidence for platy hematite grains in Sinus Meridiani, Mars. J. Geophys. Res. 107, 5126. doi:10.1029/2001JE001832

Langmuir, D., 1971. Particle size effect on the reaction Goethite = Hematite + Water. Am. J. Sci. 271, 147–156.

Larsson, D., Grönvold, K., Oskarsson, N., Gunnlaugsson, E., 2002. Hydrothermal alteration of plagioclase and growth of secondary feldspar in the Hengill Volcanic Centre, SW Iceland. J. Volcanol. Geotherm. Res. 114, 275–290. doi:10.1016/S0377-0273(01)00267-0

Le Deit, L., Flahaut, J., Quantin, C., Hauber, E., Mège, D., Bourgeois, O., Gurgurewicz, J., Massé, M., Jaumann, R., 2012. Extensive surface pedogenic alteration of the Martian Noachian crust suggested by plateau phyllosilicates around Valles Marineris. J. Geophys. Res. E Planets 117, 1–25. doi:10.1029/2011JE003983

Leroi, V., Bibring, J.-P., Berthe, M., 2009. Micromega/IR: Design and status of a near-infrared spectral microscope for in situ analysis of Mars samples. Planet. Space Sci. 57, 1068–1075. doi:10.1016/j.pss.2008.12.014

187

Leverington, D.W., 2011. A volcanic origin for the outflow channels of Mars: Key evidence and major implications. Geomorphology 132, 51–75. doi:10.1016/j.geomorph.2011.05.022

Levy, J., Head, J., Marchant, D., 2009. Thermal contraction crack polygons on Mars: Classification, distribution, and climate implications from HiRISE observations. J. Geophys. Res. 114. doi:10.1029/2008JE003273

Lewis, J.S., 1972. Metal/silicate fraccionation in the solar system. Earth Planet. Sci. Lett. 15, 286– 290.

Lodders, K., 1998. A survey of shergottite, nakhlite and chassigny meteorites whole-rock compositions. Meteorit. Planet. Sci. 33, A183–A190.

Loizeau, D., Mangold, N., Poulet, F., Bibring, J.P., Gendrin, a., Ansan, V., Gomez, C., Gondet, B., Langevin, Y., Masson, P., Neukum, G., 2007. Phyllosilicates in the Mawrth Vallis region of Mars. J. Geophys. Res. E Planets 112, 1–20. doi:10.1029/2006JE002877

Loizeau, D., Werner, S.C., Mangold, N., Bibring, J.P., Vago, J.L., 2012. Chronology of deposition and alteration in the Mawrth Vallis region, Mars. Planet. Space Sci. 72, 31–43. doi:10.1016/j.pss.2012.06.023

Lopez-Reyes, G., Rull, F., Venegas, G., Westall, F., Foucher, F., Bost, N., Sanz, A., Catalá-Espí, A., Vegas, A., Hermosilla, I., Sansano, A., Medina, J., 2013. Analysis of the scientific capabilities of the ExoMars Raman Laser Spectrometer instrument. Eur. J. Mineral. 25, 721– 733. doi:10.1127/0935-1221/2013/0025-2317

Machel, H.G., Krouse, H.R., Sassen, R., Watkin, L., 1995. Products and distinguishing criteria of bacterial and thermochemical sulfate reduction. Appl. Geochemistry 10, 373–389.

Malin, M.C., Caplinger, M.A., Edgett, K.S., Ghaemi, T.F., Ravine, M.A., Schaffner, J.A., Baker, J.M., Bardis, J.D., DiBiase, D.R., Maki, J.N., Ghaemi, F.T., Ravine, M.A., Schaffner, J.A., Baker, J.M., Bardis, J.D., DiBiase, D.R., Maki, J.N., Wilson, R.G., Bell, J.F., E., D.W., Edwards, L.J., Hallet, B., Herkenhoff, K.E., Heydari, E., Kah, L.C., Lemmon, M.T., Minitti, M.E., Olson, T.S., Parker, T.J., Rowland, S.K., Schieber, J., Sullivan, R.J., Sumner, D.Y., Thomas, P.C., Yingst, R.A., 2010. The Mars Science Laboratory (MSL) Mast-Mounted Cameras (Mastcams) flight instruments. Lunar Planet. Sci. Conf. 1123–1124.

Malin, M.C., Carr, M.H., 1999. Groundwater formation of martian valleys. Nature 397, 589–591.

Malin, M.C., Edgett, K.S., 2003. Evidence for Persistent Flow and Aqueous Sedimentation on Early Mars. Science (80-. ). 302, 1931–1935.

Malin, M.C., Edgett, K.S., 2000. Evidence for Recent Groundwater Seepage and Surface Runoff on Mars 288, 2330–2336.

Malin, M.C., Edgett, K.S., Cantor, B.A., Caplinger, M.A., Danielson, G.E., Jensen, E.H., Ravine, M.A., Sandoval, J.L., Supulver, K.D., 2010. An overview of the 1985-2006 Mars Orbiter Camera science investigation. Mars 5, 1–60. doi:10.1555/mars.2010.0001

188

Malin, M.C., Edgett, K.S., Posiolova, L. V, Mccolley, S.M., Noe Dobrea, E.Z., 2006. Present-Day Impact Cratering Rate and Contemporary Gully Activity on Mars. Science (80-. ). 314, 1573– 1578.

Manara, A., Trinchieri, G., 2011. Schiaparelli and his legacy. Mem. della Soc. Astron. Ital. 82, 209–218.

Manara, A., Wolter, A., 2011. Mars in the Schiaparelli − Lowell letters. Mem. della Soc. Astron. Ital. 82, 276–279.

Mangold, N., Allemand, P., Duval, P., Geraud, Y., Thomas, P., 2002. Experimental and theoretical deformation of ice – rock mixtures: Implications on rheology and ice content of Martian permafrost. Planet. Space Sci. 50, 385–401.

Mangold, N., Kite, E.S., Kleinhans, M.G., Newsom, H., Ansan, V., Hauber, E., Kraal, E., Quantin, C., Tanaka, K., 2012. The origin and timing of fluvial activity at Eberswalde crater, Mars. Icarus. doi:10.1016/j.icarus.2012.05.026

Marcucci, E.C., Hynek, B.M., 2014. Laboratory simulations of acid‐sulfate weathering under volcanic hydrothermal conditions: Implications for early Mars. J. Geophys. Res. Planets 199, 679–703. doi:10.1002/2013JE004439

Marcucci, E.C., Hynek, B.M., Kierein-Young, K.S., Rogers, K.L., 2013. Visible-near-infrared reflectance spectroscopy of volcanic acid-sulfate alteration in Nicaragua: Analogs for early Mars. J. Geophys. Res. Planets 118, 2213–2233. doi:10.1002/jgre.20159

Marinangeli, L., Hutchinson, I.B., Stevoli, A., Adami, G., Ambrosi, R., Amils, R., Assis Fernandes, V., Baliva, A., Basilevsky, A.T., Benedix, G., Bland, P., Böttger, A.J., Bridges, J., Caprarelli, G., Cressey, G., Critani, F., D’Alessandro, N., Delhez, R., Domeneghetti, L., Fernandez-Remolar, D., Filippone, R., Fioretti, A.M., Garcia Ruiz, J.M., Gilmore, M., Hansford, G.M., Iezzi, G., Ingley, R., Ivanov, M., Marseguerra, G., Moroz, L., Pelliciari, C., Petrinca, P., Piluso, E., Pompilio, L., Sykes, J., Westall, F., the MARS-XRD Team, 2011. The mineralogy and chemistry analyser (MARS-XRD) for the ExoMars 2018 mission, in: European Planetary Science Congress. p. EPSC-DPS2011-1232.

Marosvölgyi, K., 2009. Retrograde Alteration of Basaltic Rocks in the Þeistareykir High- Temperature Geothermal Field North-Iceland. The School for Renewable Energy Science.

Martin, W., Baross, J., Kelley, D., Russell, M.J., 2008. Hydrothermal vents and the origin of life. Nat. Rev. 6, 805–814. doi:10.1038/nrmicro1991

Marty, B., Gunnlaugsson, E., Jambon, A., Oskarsson, N., Ozima, M., Pineau, F., Torssander, P., 1991. Gas geochemistry of geothermal fluids, the Hengill area, southwest rift zone of Iceland. Chem. Geol. 91, 207–225. doi:10.1016/0009-2541(91)90001-8

Marzo, G.A., Davila, A.F., Tornabene, L.L., Dohm, J.M., Fairén, A.G., Gross, C., Kneissl, T., Bishop, J.L., Roush, T.L., McKay, C.P., 2010. Evidence for Hesperian impact-induced hydrothermalism on Mars. Icarus 208, 667–683. doi:10.1016/j.icarus.2010.03.013

189

Masursky, H., Boyce, J.M., Dial, A.L., Schaber, G.G., Strobell, M.E., 1977. Classification and Time of Formation of Martian Channels Based on Viking Data. J. Geophys. Res. 82, 4016– 4038.

Maurice, S., Diaz, E., Lopez, G., Catala, A., the RLS team, 2013. Raman Laser Spectrometer (RLS) for Exomars 2018 Rover Mission: Current Status and Science Operation Mode on Powdered Samples, in: Lunar and Planetary Science Conference. pp. 3110–3111.

Maurice, S., Wiens, R.C., Le Mouélic, S., Anderson, R., Beyssac, O., Bonal, L., Clegg, S., DeFlores, L., Dromart, G., Fischer, W., Forni, O., Gasnault, O., Grotzinger, J., Johnson, J., Martinez-Frias, J., Mangold, N., McLennan, S., Montmessin, F., Rull, F., Sharma, S., The SuperCam Team, 2015. Science Objectives of the SuperCam Instrument for th Mars 2020 rover, in: European Planetary Science Congress. p. EPSC2015-185.

McCollom, T.M., Ehlmann, B.L., Wang, A., Hynek, B.M., Moskowitz, B., Berquó, T.S., 2014. Detection of iron substitution in natroalunite-natrojarosite solid solutions and potential implications for Mars. Am. Mineral. 99, 948–964. doi:10.2138/am.2014.4617

McCollom, T.M., Hynek, B.M., Rogers, K., Moskowitz, B., Berquó, T.S., 2013a. Chemical and mineralogical trends during acid-sulfate alteration of pyroclastic basalt at Cerro Negro volcano and implications for early Mars. J. Geophys. Res. Planets 118, 1719–1751. doi:10.1002/jgre.20114

McCollom, T.M., Robbins, M., Moskowitz, B., Berquõ, T.S., Jöns, N., Hynek, B.M., 2013b. Experimental study of acid-sulfate alteration of basalt and implications for sulfate deposits on Mars. J. Geophys. Res. Planets 118, 577–614. doi:10.1002/jgre.20044

McCubbin, F.M., Tosca, N.J., Smirnov, A., Nekvasil, H., Steele, A., Fries, M., Lindsley, D.H., 2009. Hydrothermal jarosite and hematite in a pyroxene-hosted melt inclusion in martian meteorite Miller Range (MIL) 03346: Implications for magmatic-hydrothermal fluids on Mars. Geochim. Cosmochim. Acta 73, 4907–4917. doi:10.1016/j.gca.2009.05.031

McCusker, L.B., Von Dreele, R.B., Cox, D.E., Louer, D., Scardi, P., 1999. Rietveld refinement guidelines. J. Appl. Crystallogr. 32, 36–50.

Mchenry, L.J., 2009. Element mobility during zeolitic and argillic alteration of volcanic ash in a closed-basin lacustrine environment: Case study Olduvai Gorge, Tanzania. Chem. Geol. 265, 540–552. doi:10.1016/j.chemgeo.2009.05.019

McHenry, L.J., Carson, G.L., Hynek, B.M., 2016. Assessing Icelandic Fumaroles and Hot Springs as Analogs for Mars, in: Lunar and Planetary Science Conference. pp. 1207–1208.

McHenry, L.J., Chevrier, V., Schörder, C., 2011. Jarosite in a Pleistocene East African saline- alkaline paleolacustrine deposit: Implications for Mars aqueous geochemistry. J. Geophys. Res. E Planets 116. doi:10.1029/2010JE003680

McKeown, N.K., Bishop, J.L., Noe Dobrea, E.Z., Ehlmann, B.L., Parente, M., Mustard, J.F., Murchie, S.L., Swayze, G. a., Bibring, J.P., Silver, E. a., 2009. Characterization of

190

phyllosilicates observed in the central Mawrth Vallis region, Mars, their potential formational processes, and implications for past climate. J. Geophys. Res. 114. doi:10.1029/2008JE003301

McSween, H.Y., Murchie, S.L., Crisp, J. a., Bridges, N.T., Anderson, R.C., Bell, J.F., Britt, D.T., Briickner, J., Dreibus, G., Economou, T., Ghosh, a., Golombek, M.P., Greenwood, J.P., Johnson, J.R., Moore, H.J., Morris, R. V., Parker, T.J., Rieder, R., Singer, R., H., W., 1999. Chemical, multispectral, and textural constraints on the composition and origin of rocks at the Mars Pathfinder landing site, Journal of Geophysical Research. doi:10.1029/98je02551

McSween, H.Y., Ruff, S.W., Morris, R. V, Bell, J.F., Herkenhoff, K.E., Gellert, R., the ATHENA Science Team, Branch, A., Survey, G., 2006. Backstay and Irvine: Alkaline volcanic rocks from Gusev crater, Mars, in: Lunar and Planetary Science Conference. pp. 1120–1121.

McSween, H.Y., Taylor, G.J., Wyatt, M.B., 2009. Elemental composition of the Martian crust. Science (80-. ). 324, 736–739. doi:10.1126/science.1165871

McSween, H.Y., Wyatt, M.B., Gellert, R., Bell, J.F., Morris, R. V., Herkenhoff, K.E., Crumpler, L.S., Milam, K. a., Stockstill, K.R., Tornabene, L.L., Arvidson, R.E., Bartlett, P., Blaney, D., Cabrol, N. a., Christensen, P.R., Clark, B.C., Crisp, J. a., Des Marais, D.J., Economou, T., Farmer, J.D., Farrand, W., Ghosh, A., Golombek, M., Gorevan, S., Greeley, R., Hamilton, V.E., Johnson, J.R., Joliff, B.L., Klingelhöfer, G., Knudson, a. T., McLennan, S., Ming, D., Moersch, J.E., Rieder, R., Ruff, S.W., Schröder, C., de Souza, P. a., Squyres, S.W., Wänke, H., Wang, A., Yen, A., Zipfel, J., 2006. Characterization and petrologic interpretation of olivine-rich basalts at Gusev Crater, Mars. J. Geophys. Res. 111. doi:10.1029/2005JE002477

Meslin, P., Gasnault, O., Forni, O., Schröder, S., Cousin, A., Berger, G., Clegg, S.M., Lasue, J., Maurice, S., Sautter, V., Le Mouélic, S., Wiens, R.C., Fabre, C., Goetz, W., Bish, D., Mangold, N., Ehlmann, B., Lanza, N., Harri, A.-M., Anderson, R., Rampe, E., McConnochie, T.H., Pinet, P., Blaney, D., Léveillé, R., Archer, D., Barraclough, B., Bender, S., Blake, D., Blank, J.G., Bridges, N., Clark, B.C., DeFlores, L., Delapp, D., Dromart, G., Dyar, M.D., Fisk, M., Gondet, B., Grotzinger, J., Herkenhoff, K., Johnson, J., Lacour, J.-L., Langevin, Y., Leshin, L., Lewin, E., Madsen, M.B., Melikechi, N., Mezzacappa, A., Mischna, M.A., Moores, J.E., Newsom, H., Ollila, A., Perez, R., Renno, N., Sirven, J.-B., Tokar, R., de la Torre, M., D’Uston, L., Vaniman, D., Yingst, A., MSL Science Team, 2013. Soil Diversity and Hydration at Gale Crater, Mars. Science (80-. ). 341. doi:10.1126/science.1238670

Michalski, J., Poulet, F., Bibring, J.P., Mangold, N., 2010. Analysis of phyllosilicate deposits in the Nili Fossae region of Mars: Comparison of TES and OMEGA data. Icarus 206, 269–289. doi:10.1016/j.icarus.2009.09.006

Michalski, J.R., Bleacher, J.E., 2013. Supervolcanoes within an ancient volcanic province in Arabia Terra , Mars. Nature 502, 47–52. doi:10.1038/nature12482

Michalski, J.R., Kraft, M.D., Diedrich, T., Sharp, T.G., Christensen, P.R., 2003. Thermal emission spectroscopy of the silica polymorphs and considerations for remote sensing of Mars. Geophys. Res. Lett. 30. doi:10.1029/2003GL018354

191

Michalski, J.R., Kraft, M.D., Sharp, T.G., Williams, L.B., Christensen, P.R., 2005. Mineralogical constraints on the high-silica martian surface component observed by TES. Icarus 174, 161– 177. doi:10.1016/j.icarus.2004.10.022

Michalski, J.R., Niles, P.B., Cuadros, J., Baldridge, a. M., 2013. Multiple working hypotheses for the formation of compositional stratigraphy on Mars: Insights from the Mawrth Vallis region. Icarus 226, 816–840. doi:10.1016/j.icarus.2013.05.024

Milliken, R.E., Hurowitz, J.A., Bish, D.L., Grotzinger, J.P., Wiens, R., 2016. The chemical and mineralogical stratigraphy of lower Mt. Sharp: Relating rover observations to orbital predictions, in: Lunar and Planetary Science Conference. pp. 1495–1496.

Milliken, R.E., Swayze, G.A., Arvidson, R.E., Bishop, J.L., Clark, R.N., Ehlmann, B.L., Green, R.O., Grotzinger, J.P., 2008. Opaline silica in young deposits on Mars 847–850. doi:10.1130/G24967A.1

Ming, D.W., Gellert, R., Morris, R. V., Arvidson, R.E., Brückner, J., Clark, B.C., Cohen, B.A., D’Uston, C., Economou, T., Fleischer, I., Klingelhöfer, G., McCoy, T.J., Mittlefehldt, D.W., Schmidt, M.E., Schröder, C., Squyres, S.W., Tréguier, E., Yen, A.S., Zipfel, J., 2008. Geochemical properties of rocks and soils in Gusev Crater, Mars: Results of the Alpha Particle X-Ray Spectrometer from Cumberland Ridge to Home Plate. J. Geophys. Res. 113. doi:10.1029/2008JE003195

Mínguez, H.A., Ortega, L., Lunar, R., Martínez-Frías, J., Piña, R., 2011. Mineralogy of the Hydrothermal Alteration in the Námafjall Geothermal Field (Iceland). Rev. la Soc. Española Mineral. 15, 25–26.

Minitti, M.E., Weitz, C.M., Lane, M.D., Bishop, J.L., 2007. Morphology, chemistry, and spectral properties of Hawaiian rock coatings and implications for Mars. J. Geophys. Res. 112. doi:10.1029/2006JE002839

Monders, A.G., Medard, E., Grove, T.L., 2007. Phase equilibrium investigations of the Adirondack class basalts from the Gusev plains, Gusev crater, Mars. Meteorit. Planet. Sci. 42, 131–148. doi:10.1111/j.1945-5100.2007.tb00222.x

Montgomery, D.R., Gillespie, A., 2005. Formation of Martian outflow channels by catastrophic dewatering of evaporite deposits. Geology 33, 625–628. doi:10.1130/G21270.1

Morris, R. V., Klingelhöfer, G., Schröder, C., Rodionov, D.S., Yen, A.S., Ming, D.W., de Souza Jr., P.A., Wdowiak, T., Fleischer, I., Gellert, R., Bernhardt, B., Bonnes, U., Cohen, B.A., Evlanov, E.N., Foh, J., Gütlich, P., Kankeleit, E., McCoy, T., Mittlefehldt, D.W., Renz, F., Schmidt, M.E., Zubkov, B., Squyres, S.W., Arvidson, R.E., 2006. Mössbauer mineralogy of rock, soil, and dust at Meridiani Planum, Mars: Opportunity’s journey across sulfate-rich outcrop, basaltic sand and dust, and hematite lag deposits. J. Geophys. Res. 111. doi:10.1029/2006JE002791

Morris, R. V., Klingelhöfer, G., Schröer, C., Fleischer, I., Ming, D.W., Yen, A.S., Gellert, R., Arvidson, R.E., Rodionov, D.S., Crumpler, L.S., Clark, B.C., Cohen, B.A., McCoy, T.J.,

192

Mittlefehldt, D.W., Schmidt, M.E., de Souza, P.A., Squyres, S.W., 2008. Iron mineralogy and aqueous alteration from Husband Hill through Home Plate at Gusev Crater, Mars: Results from the Mössbauer instrument on the Spirit Mars Exploration Rover. J. Geophys. Res. 113. doi:10.1029/2008JE003201

Morris, R. V., Ming, D.W., Graff, T.G., Arvidson, R.E., Bell, J.F., Squyres, S.W., Mertzman, S.A., Gruener, J.E., Golden, D.C., Le, L., Robinson, G.A., 2005. Hematite spherules in basaltic tephra altered under aqueous, acid-sulfate conditions on Mauna Kea volcano, Hawaii: Possible clues for the occurrence of hematite-rich spherules in the Burns formation at Meridiani Planum, Mars. Earth Planet. Sci. Lett. 240, 168–178. doi:10.1016/j.epsl.2005.09.044

Morris, R. V., Vaniman, D.T., Blake, D.F., Gellert, R., Chipera, S.J., Rampe, E.B., Ming, D.W., Morrison, S.M., Downs, R.T., Treiman, A.H., Yen, A.S., Grotzinger, J.P., Achilles, C.N., Bristow, T.F., Crisp, J.A., Des Marais, D.J., Farmer, J.D., Fendrich, K. V., Frydenvang, J., Graff, T.G., Morookian, J.-M., Stolper, E.M., Schwenzer, S.P., 2016. Silicic volcanism on Mars evidenced by tridymite in high-SiO2 sedimentary rock at Gale crater. Proc. Natl. Acad. Sci. doi:10.1073/pnas.1607098113

Morris, R. V, Golden, D.C., Bell, J.F., Shelfer, T.D., Scheinost, A.C., Hinman, N.W., Furniss, G., Mertzman, S.A., Bishop, J.L., Ming, D.W., Allen, C.C., Britt, D.T., 2000. Mineralogy, composition, and alteration of Mars Pathfinder rocks and soils: Evidence from multispectral, elemental, and magnetic data on terrestrial analogue, SNC meteorite, and Pathfinder samples. J. Geophys. Res. 105, 1757–1817. doi:10.1029/1999JE001059

Mouginis-Mark, P.J., Rowland, S.K., 2008. Lava flows at Arsia Mons, Mars: Insights from a graben imaged by HiRISE. Icarus 198, 27–36. doi:10.1016/j.icarus.2008.06.015

Moune, S., Sigmarsson, O., Thordarson, T., Gauthier, P., 2007. Recent volatile evolution in the magmatic system of Hekla volcano, Iceland 255, 373–389. doi:10.1016/j.epsl.2006.12.024

Murchie, S., Arvidson, R., Bedini, P., Beisser, K., Bibring, J.P., Bishop, J., Boldt, J., Cavender, P., Choo, T., Clancy, R.T., Darlington, E.H., Des Marais, D., Espiritu, R., Fort, D., Green, R., Guinness, E., Hayes, J., Hash, C., Heffernan, K., Hemmler, J., Heyler, G., Humm, D., Hutcheson, J., Izenberg, N., Lee, R., Lees, J., Lohr, D., Malaret, E., Martin, T., McGovern, J. a., McGuire, P., Morris, R., Mustard, J., Pelkey, S., Rhodes, E., Robinson, M., Roush, T., Schaefer, E., Seagrave, G., Seelos, F., Silverglate, P., Slavney, S., Smith, M., Shyong, W.J., Strohbehn, K., Taylor, H., Thompson, P., Tossman, B., Wirzburger, M., Wolff, M., 2007. Compact Connaissance Imaging Spectrometer for Mars (CRISM) on Mars Reconnaissance Orbiter (MRO). J. Geophys. Res. E Planets 112, 1–57. doi:10.1029/2006JE002682

Murchie, S., Roach, L., Seelos, F., Milliken, R., Mustard, J., Arvidson, R., Wiseman, S., Lichtenberg, K., Andrews-hanna, J., Bishop, J., Bibring, J., Parente, M., Morris, R., 2009. Evidence for the origin of layered deposits in Candor Chasma, Mars, from mineral composition and hydrologic modeling. J. Geophys. Res. 114. doi:10.1029/2009JE003343

Murchie, S.L., Mustard, J.F., Ehlmann, B.L., Milliken, R.E., Bishop, J.L., McKeown, N.K., Noe Dobrea, E.Z., Seelos, F.P., Buczkowski, D.L., Wiseman, S.M., Arvidson, R.E., Wray, J.J., 193

Swayze, G., Clark, R.N., Des Marais, D.J., McEwen, A.S., Bibring, J.-P., 2009a. A synthesis of Martian aqueous mineralogy after 1 Mars year of observations from the Mars Reconnaissance Orbiter. J. Geophys. Res. 114. doi:10.1029/2009JE003342

Murchie, S.L., Seelos, F.P., Hash, C.D., Humm, D.C., Malaret, E., McGovern, J.A., Choo, T.H., Seelos, K.D., Buczkowski, D.L., Morgan, M.F., Barnouin-Jha, O.S., Nair, H., Taylor, H.W., Patterson, G.W., Harvel, C. a., Mustard, J.F., Arvidson, R.E., McGuire, P., Smith, M.D., Wolff, M.J., Titus, T.N., Bibring, J.P., Poulet, F., 2009b. Compact Reconnaissance Imaging Spectrometer for Mars investigation and data set from the Mars Reconnaissance Orbiter’s primary science phase. J. Geophys. Res. 114. doi:10.1029/2009JE003344

Musselwhite, D.S., Swindle, T.D., Lunine, J.I., 2001. Liquid CO2 Breakout and the Formation of Recent Small Gullies on Mars. Geophys. Res. Lett. 28, 1283–1285.

Mustard, J.F., Murchie, S.L., Pelkey, S.M., Ehlmann, B.L., Milliken, R.E., Grant, J. a, Bibring, J.- P., Poulet, F., Bishop, J., Dobrea, E.N., Roach, L., Seelos, F., Arvidson, R.E., Wiseman, S., Green, R., Hash, C., Humm, D., Malaret, E., McGovern, J. a, Seelos, K., Clancy, T., Clark, R., Marais, D.D., Izenberg, N., Knudson, a, Langevin, Y., Martin, T., McGuire, P., Morris, R., Robinson, M., Roush, T., Smith, M., Swayze, G., Taylor, H., Titus, T., Wolff, M., 2008. Hydrated silicate minerals on Mars observed by the Mars Reconnaissance Orbiter CRISM instrument. Nature 454, 305–309. doi:10.1038/nature07097

Mustard, J.F., Poulet, F., Gendrin, A., Bibring, J.-P., Langevin, Y., Gondet, B., Mangold, N., Bellucci, G., Altieri, F., 2005. Olivine and pyroxene diversity in the crust of Mars. Science 307, 1594–7. doi:10.1126/science.1109098

Nedell, S.S., Squyres, S.W., Andersen, D.W., 1987. Origin and Evolution of the Layered Deposits in the Valles Marineris , Mars outflow channels in Chryse Basin and Lu-. Icarus 70, 409–441.

Neukum, G., Jaumann, R., Hoffmann, H., Hauber, E., Head, J.W., Basilevsky, A.T., Ivanov, B.A., Werner, S.C., van Gasselt, S., Murray, J.B., McCord, T., The HRSC Co-Investigator Team, 2004. Recent and episodic volcanic and glacial activity on Mars revealed by the High Resolution Stereo Camera. Nature 432.

Nicholson, H., Latin, D., 1992. Olivine Tholeiites from Krafla, Iceland: Evidence for Variations in Melt Fraction within a Plume. J. Petrol. 33, 1105–1124.

Novák, Z., 2006. El área volcánica de Momotombo (Nicaragua). Krystalinikum 31, 25–46.

Nyquist, L.E., Bogard, D.D., Shih, C.Y., Greshake, A., Stoffler, D., Eugster, O., 2001. Ages and geologic histories of Martian meteorites, in: Chronology and Evolution of Mars. Kluwer Academic Publishers, pp. 105–164. doi:10.1023/A:1011993105172

Okon, A.B., 2010. Mars Science Laboratory Drill, in: Proceedings of the 40th Aerospace Mechanisms Symposium.

Olympus Scientific Solutions, 2014. TERRA — X-Ray Diffraction and X-Ray Fluorescence Analyzer: User’s Manual.

194

Óskarsson, F., Ármannsson, H., Ólafsson, M., Sveinbjörnsdóttir, Á.E., Markússon, S.H., 2013. The Theistareykir geothermal field, NE Iceland: fluid chemistry and production properties. Procedia Earth Planet. Sci. 7, 644–647. doi:10.1016/j.proeps.2013.03.185

Paul, A., Zaman, M.S., 1978. The relative influences of AI203 and Fe203 on the chemical durability of silicate glasses at different pH values. J. Mater. Sci. 13, 1499–1502.

PDS Geosciences Node, CRISM Spectral Library Working Group, 2014. CRISM Spectral Library. Washingt. Univ. St. Louis.

Peretyazhko, T.S., Niles, P.B., Sutter, B., Morris, R. V, Agresti, D.G., Le, L., Ming, D.W., 2018. Smectite formation in the presence of sulfuric acid Implications for acidic smectite formation on early Mars. Geochim. Cosmochim. Acta 220, 248–260. doi:10.1016/j.gca.2017.10.004

Peterson, C., 1981. A secondary origin for the central plateay of Hebes Chasma, in: Lunar and Planetary Science Conference. pp. 1459–1471.

Phillips, R.J., Zuber, M.T., Solomon, S.C., Golombek, M.P., Jakosky, B.M., Banerdt, W.B., Smith, D.E., Williams, R.M., Hynek, B.M., Aharonson, O., Hauck, S. a, 2001. Ancient geodynamics and global-scale hydrology on Mars. Science 291, 2587–2591.

Pirajno, F., 2010. Hydrothermal Processes and Mineral Systems. Springer Science+Business Media B.V.

Plumb, R.C., Tantayanon, R., Libby, M., Xu, W.W., 1989. Chemical model for Viking biology experiments: implications for the composition of the martian regolith. Lett. to Nat. 338, 633– 635.

Pollack, J.B., Kasting, J.F., Richardson, S.M., Poliakoff, K., 1987. The Case for a Wet, Warm Climate on Early Mars. Icarus 71, 203–224.

Pondrelli, M., Pio Rossi, A., Marinangeli, L., Hauber, E., Gwinner, K., Baliva, A., Di Lorenzo, S., 2008. Evolution and depositional environments of the Eberswalde fan delta, Mars. Icarus 197, 429–451. doi:10.1016/j.icarus.2008.05.018

Poulet, F., Carter, J., Bishop, J.L., Loizeau, D., Murchie, S.M., 2014. Mineral abundances at the final four curiosity study sites and implications for their formation. Icarus 231, 65–76. doi:10.1016/j.icarus.2013.11.023

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

Preston, L.J., Dartnell, L.R., 2014. Planetary habitability: lessons learned from terrestrial analogues. Int. J. Astrobiol. 13, 81–98. doi:10.1017/S1473550413000396

Prosser, J.T., Carr, M.J., 1987. Poás volcano, Costa Rica: Geology of the summit region and spatial

195

and temporal variations among the most recent lavas. J. Volcanol. Geotherm. Res. 33, 131– 146. doi:10.1016/0377-0273(87)90057-6

Quantin, C., Carter, J., Thollot, P., Broyer, J., Lozach, L., Davis, J., Grindrod, P., Pajola, M., Baratti, E., Rossato, S., Allemand, P., Bultel, B., Leyrat, C., Fernando, J., Ody, A., 2016. Oxia Planum, the landing site for ExoMars 2018, in: Lunar and Planetary Science Conference. pp. 2863–2864.

Rampe, E.B., Ming, D.W., Morris, R. V., Blake, D.F., Bristow, T.F., Chipera, S.J., Vaniman, D.T., Yen, A.S., Grotzinger, J.P., Downs, R.T., Morrison, S.M., Peretyazhko, T., Achilles, C.N., Bish, D.L., Cavanagh, P.D., Craig, P.I., Crisp, J.A., Fairén, A.G., Marais, D.J. Des, Farmer, J.D., Fendrich, K. V., Morookian, J.M., Treiman, A.H., 2016. Diagenesis in the Murray Formation, Gale Crater, Mars, in: Lunar and Planetary Science Conference. pp. 2543–2544. doi:10.1038/ngeo2474.

Rampe, E.B., Ming, D.W., Vaniman, D.T., Blake, D.F., Chipera, S., Morris, R. V., Bish, D.L., Cavanagh, P., Achilles, C., Bristow, T., Fairen, A., Morrison, S.M., Treiman, A.H., Crisp, J.A., Downs, R.T., Farmer, J.D., Fendrich, K., Morookian, J., Affiliation:, 2015. Evidence for Acid-Sulfate Alteration in the Pahrump Hills Region, Gale Crater, Mars, in: American Geophysical Union, Fall Meeting. p. abstract #P53F-04.

Rampe, E.B., Morris, R. V., Ming, D.W., Archer, P.D., Bish, D.L., Chipera, S.J., Vaniman, D.T., Blake, D.F., Bristow, T.F., Sutter, B., Farmer, J.D., Downs, R.T., Leveille, R., Achilles, C.A., Crisp, J.A., Des Marais, D.J., Morookian, J.M., Morrison, S.M., Sarrazin, P.C., Spanovich, N., Treiman, A.H., Yen, A.S., 2014. Characterizing the phyllosilicate component of the Sheepbed mudstone in Gale Crater, Mars using laboratory XRD and EGA. 45th Lunar Planet. Sci. Conf. 2, 2–3. doi:10.1126/science.1243480.

Rathbun, J.A., Squyres, S.W., 2002. Hydrothermal Systems Associated with Martian Impact Craters. Icarus 157, 362–372. doi:10.1006/icar.2002.6838

Reagan, M.K., Gill, J.B., 1989. Coexisting calcalkaline and high-niobium basalts from Turrialba Volcano, Costa Rica: Implications for residual titanates in arc magma sources. J. Geophys. Res. 94, 4619–4633.

Reed, S.J.B., 2005. Electron Microprobe Analysis and Scanning Electron Microscopy in Geology. Cambridge University Press.

Reysenbach, A.-L., Cady, S.L., 2001. Microbiology of ancient and modern hydrothermal systems. Trends Microbiol. 9, 79–86.

Rice, M.S., Bell, J.F., Cloutis, E. a., Wang, A., Ruff, S.W., Craig, M. a., Bailey, D.T., Johnson, J.R., de Souza, P. a., Farrand, W.H., 2010. Silica-rich deposits and hydrated minerals at Gusev Crater, Mars: Vis-NIR spectral characterization and regional mapping. Icarus 205, 375–395. doi:10.1016/j.icarus.2009.03.035

Rice, M.S., Cloutis, E.A., Bell, J.F., Bish, D.L., Horgan, B.H., Mertzman, S.A., Craig, M.A., Renaut, R.W., Gautason, B., Mountain, B., 2013. Reflectance spectra diversity of silica-rich

196

materials: Sensitivity to environment and implications for detections on Mars. Icarus 223, 499–533. doi:10.1016/j.icarus.2012.09.021

Rickard, D., 1997. Kinetics of pyrite formation by the H2S oxidation of iron (II) monosulfide in aqueous solutions between 25 and 125° C: The rate equation. Geochim. Cosmochim. Acta 61, 115–134.

Rickard, D., Luther, G.W., 1997. Pergamon Kinetics of pyrite formation by the H2S oxidation of iron (II) monosulfide in aqueous solutions between 25 and 125° C: The mechanism. Geochim. Cosmochim. Acta 61, 135–147.

Rieder, R., Gellert, R., Anderson, R.C., Brückner, J., Clark, B.C., Dreibus, G., Economou, T., Klingelhöfer, G., Lugmair, G.W., Ming, D.W., Squyres, S.W., D’Uston, C., Wänke, H., Yen, A., Zipfel, J., 2004. Chemistry of rocks and soils at Meridiani Planum from the Alpha Particle X-ray Spectrometer. Science (80-. ). 306, 1746–1749. doi:10.1126/science.1104358

Rieder, R., Gellert, R., Brückner, J., Klingelhöfer, G., Dreibus, G., Yen, A., Squyres, S.W., 2003. The new Athena alpha particle X-ray spectrometer for the Mars Exploration Rovers. J. Geophys. Res. 108, 8066. doi:10.1029/2003JE002150

Rieder, R., Wänke, H., Economou, T., Turkevich, A., 1997. Determination of the chemical composition of Martian soil and rocks: The alpha proton X ray spectrometer. J. Geophys. Res. 102, 4027–4044.

Righter, K., Pando, K., Danielson, L.R., 2009. Experimental evidence for sulfur-rich martian magmas: Implications for volcanism and surficial sulfur sources. Earth Planet. Sci. Lett. 288, 235–243. doi:10.1016/j.epsl.2009.09.027

Robbins, S.J., Di Achille, G., Hynek, B.M., 2011. The volcanic history of Mars: High-resolution crater-based studies of the calderas of 20 volcanoes. Icarus 211, 1179–1203. doi:10.1016/j.icarus.2010.11.012

Rodríguez, A., van Bergen, M.J., 2017. Superficial alteration mineralogy in active volcanic systems: An example of Poás volcano, Costa Rica. J. Volcanol. Geotherm. Res. doi:10.1016/j.jvolgeores.2017.04.006

Rodríguez, A., van Bergen, M.J., 2015. Volcanic hydrothermal systems as potential analogues of Martian sulphate-rich terrains. Netherlands J. Geosci. April, 1–17. doi:10.1017/njg.2015.12

Rogers, A.D., Hamilton, V.E., 2015. Compositional provinces of Mars from statistical analyses of TES, GRS, OMEGA and CRISM data. J. Geophys. Res. Planets 120, 62–91. doi:10.1002/2014JE004690.Received

Roggensack, K., Herving, R.L., McKnight, S.B., Williams, S.N., 1997. Explosive Basalt Volcanism from Cerro Negro Volcano: Influence of Volatiles on Eruptive Style. Science (80- . ). 277, 1639–1642.

Rubin, A.E., Warren, P.H., Greenwood, J.P., Verish, R.S., Leshin, L.A., Hervig, R.L., Clayton,

197

R.N., Mayeda, T.K., 2000. Los Angeles: The most differentiated basaltic martian meteorite. Geology 28, 1011–1014. doi:10.1130/0091-7613(2000)28<1011:LATMDB>2.0.CO

Ruff, S.W., 2015. New Observations Reveal a Former Hot Spring Environment With High Habitability and Preservation Potential in Gusev Crater, Mars, in: Lunar and Planetary Science Conference. pp. 1613–1614.

Ruff, S.W., Christensen, P.R., 2002. Bright and dark regions on Mars: Particle size and mineralogical characteristics based on Thermal Emission Spectrometer data. J. Geophys. Res. 107. doi:10.1029/2001JE001580

Ruff, S.W., Farmer, J.D., Calvin, W.M., Herkenhoff, K.E., Johnson, J.R., Morris, R. V., Rice, M.S., Arvidson, R.E., Bell, J.F., Christensen, P.R., Squyres, S.W., 2011. Characteristics, distribution, origin, and significance of opaline silica observed by the Spirit rover in Gusev crater, Mars. J. Geophys. Res. 116. doi:10.1029/2010JE003767

Rull, F., Sansano, A., Díaz, E., Canora, C.P., Moral, A.G., Tato, C., Colombo, M., Belenguer, T., Fernández, M., Manfredi, J.A.R., Canchal, R., Dávila, B., Jiménez, A., Gallego, P., Ibarmia, S., Prieto, J.A.R., Santiago, A., Pla, J., Ramos, G., Díaz, C., González, C., 2011. ExoMars Raman Laser Spectrometer for Exomars, in: SPIE Optical Engineering + Applications. p. 81520J–81520J. doi:10.1117/12.896787

Saemundsson, K., 1991. Geology of the Krafla System. Nattura Myvatns 24–95.

Sæmundsson, K., 2007. The Geology of Theistareykir (in Icelandic). Reykjavik, Iceland.

Sarrazin, P., Blake, D., Feldman, S., Chipera, S., Vaniman, D., Bish, D., 2005. Field deployment of a portable X-ray diffraction/X-ray flourescence instrument on Mars analog terrain. Powder Diffr. 20, 128–133. doi:10.1154/1.1913719

Sautter, V., Toplis, M.J., Beck, P., Mangold, N., Wiens, R., Pinet, P., Cousin, A., Maurice, S., LeDeit, L., Hewins, R., Gasnault, O., Quantin, C., Forni, O., Newsom, H., Meslin, P.Y., Wray, J., Bridges, N., Payré, V., Rapin, W., Le Mouélic, S., 2016. Magmatic complexity on early Mars as seen through a combination of orbital, in-situ and meteorite data. Lithos 254– 255, 36–52. doi:10.1016/j.lithos.2016.02.023

Sautter, V., Toplis, M.J., Wiens, R.C., Cousin, A., Fabre, C., Gasnault, O., Maurice, S., Forni, O., Lasue, J., Ollila, A., Bridges, J.C., Mangold, N., Le Mouélic, S., Fisk, M., Meslin, P.-Y., Beck, P., Pinet, P., Le Deit, L., Rapin, W., Stolper, E.M., Newsom, H., Dyar, D., Lanza, N., Vaniman, D., Clegg, S., Wray, J.J., 2015. In situ evidence for continental crust on early Mars. Nat. Geosci. 8, 605–609. doi:10.1038/ngeo2474

Schmidt, M.E., Farrand, W.H., Johnson, J.R., Schröder, C., Hurowitz, J.A., McCoy, T.J., Ruff, S.W., Arvidson, R.E., Des Marais, D.J., Lewis, K.W., Ming, D.W., Squyres, S.W., de Souza, P.A., 2009. Spectral, mineralogical, and geochemical variations across Home Plate, Gusev Crater, Mars indicate high and low temperature alteration. Earth Planet. Sci. Lett. 281, 258– 266. doi:10.1016/j.epsl.2009.02.030

198

Schmidt, M.E., Ruff, S.W., McCoy, T.J., Farrand, W.H., Johnson, J.R., Gellert, R., Ming, D.W., Morris, R. V., Cabrol, N., Lewis, K.W., Schroeder, C., 2008. Hydrothermal origin of halogens at Home Plate, Gusev Crater. J. Geophys. Res. 113. doi:10.1029/2007JE003027

Schröder, C., Schwenzer, S.P., 2017. Evidence for impact-induced hydrothermal clay mineral formation at Endeavour Crater, Mars, in: Lunar and Planetary Science Conference. pp. 2429– 2430. doi:10.1126/science.1248097.

Schuessler, J.A., Schoenberg, R., Sigmarsson, O., 2009. Iron and lithium isotope systematics of the Hekla volcano, Iceland — Evidence for Fe isotope fractionation during magma differentiation. Chem. Geol. 258, 78–91. doi:10.1016/j.chemgeo.2008.06.021

Schultz, R.A., Tanaka, K.L., 1994. Lithospheric-scale buckling and thrust structures on Mars: The Coprates rise and south Tharsis ridge belt. J. Geophys. Res. 99, 8371–8385.

Schulze-Makuch, D., Dohm, J.M., Fan, C., Fairén, A.G., Rodriguez, J. a P., Baker, V.R., Fink, W., 2007. Exploration of hydrothermal targets on Mars. Icarus 189, 308–324. doi:10.1016/j.icarus.2007.02.007

Schwenzer, S.P., Kring, D.A., 2013. Alteration minerals in impact-generated hydrothermal systems - Exploring host rock variability. Icarus 226, 487–496. doi:10.1016/j.icarus.2013.06.003

Schwenzer, S.P., Kring, D.A., 2009. Impact-generated hydrothermal systems capable of forming phyllosilicates on Noachian Mars. Geology 37, 1091–1094. doi:10.1130/G30340A.1

Scott, D.H., Tanaka, K.L., 1986. Geologic Map of the Western Equatorial Region of Mars.

Sharp, R.P., 1973. Mars: Troughed Terrain. J. Geophys. Res. 78, 4063–4072.

Sharp, R.P., Malin, M.C., 1975. Channels on Mars. Geol. Soc. Am. Bull. 86, 593–609.

Singer, R.B., 1982. Spectral Evidence for the Mineralogy of High-Albedo Soils and Dust on Mars. J. Geophys. Res. 87, 10159–10168. doi:10.1029/JB087iB12p10159

Sinton, J., Gronvold, K., Saemundsson, K., 2005. Postglacial eruptive history of the Western Volcanic Zone, Iceland. Geochemistry Geophys. Geosystems 6. doi:Q12009\r10.1029/2005gc001021

Siva, V., Sinha, R.K., 2017. Amazonian fluvial outflow channels in Jovis Tholus region, Mars. J. Geophys. Res. Planets. doi:10.1002/2016JE005237

Skok, J.R., Mustard, J.F., Ehlmann, B.L., Milliken, R.E., Murchie, S.L., 2010. Silica deposits in the Nili Patera caldera on the Syrtis Major volcanic complex on Mars. Nat. Geosci. 3, 838– 841. doi:10.1038/ngeo990

Smith, D.E., Zuber, M.T., Frey, H. V., Garvin, J.B., Head, James, W., Muhleman, D.O., Pettengill, G.H., Phillips, R.J., Solomon, S.C., Zwally, H.J., Banerdt, W.B., Duxbury, T.C., Golombek,

199

M.P., Lemoine, F.G., Neumann, G.A., Rowlands, D.D., Aharonson, O., Ford, P.G., Ivanov, A.B., Johnson, C.L., Mcgovern, P.J., Abshire, J.B., Afzal, R.S., Sun, X., 2001. Mars Orbiter Laser Altimeter: Experiment summary after the first year of global mapping of Mars. J. Geophys. Res. 106, 23689–23722.

Smith, M.R., Bandfield, J.L., 2012. Geology of quartz and hydrated silica-bearing deposits near Antoniadi Crater, Mars. J. Geophys. Res. 117. doi:10.1029/2011JE004038

Smith, P.H., Bell, J.F., Bridges, N.T., Britt, D.T., Gaddis, L., Greeley, R., Keller, H.U., Herkenhoff, K.E., Jaumann, R., Johnson, J.R., Kirk, R.L., Lemmon, M., Maki, J.N., Malin, M.C., Murchie, S.L., Oberst, J., Parker, T.J., Reid, R.J., Sablotny, R., Soderblom, L.A., Stoker, C., Sullivan, R., Thomas, N., Tomasko, M.G., Ward, W., Wegryn, E., 1997. Results from the Mars Pathfinder Camera. Science (80-. ). 278, 1758–1765.

Squyres, S.W., Aharonson, O., Clark, B.C., Cohen, B.A., Crumpler, L., de Souza, P.A., Farrand, W.H., Gellert, R., Grant, J., Grotzinger, J.P., Haldemann, A.F.C., Johnson, J.R., Klingelhöfer, G., Lewis, K.W., Li, R., McCoy, T., McEwen, A.S., McSween, H.Y., Ming, D.W., Moore, J.M., Morris, R. V., Parker, T.J., Rice, J.W., Ruff, S., Schmidt, M., Schröder, C., Soderblom, L.A., Yen, A., 2007. Pyroclastic Activity at Home Plate in Gusev Crater, Mars. Science (80- . ). 316, 738–742. doi:10.1126/science.1139045

Squyres, S.W., Arvidson, R.E., Ruff, S., Gellert, R., Morris, R. V., Ming, D.W., Crumpler, L., Farmer, J.D., Des Marais, D.J., Yen, A., McLennan, S.M., Calvin, W., Bell, J.F., Clark, B.C., Wang, A., McCoy, T.J., Schmidt, M.E., de Souza, P.A., 2008. Detection of Silica-Rich Deposits on Mars. Science (80-. ). 320, 1063–1068. doi:10.1126/science.1155429

Squyres, S.W., Kasting, J.F., 1994. Early Mars: How Warm and How Wet? Science (80-. ). 265, 744–749.

Squyres, S.W., Knoll, A.H., 2005. Sedimentary rocks at Meridiani Planum: Origin, diagenesis, and implications for life on Mars. Earth Planet. Sci. Lett. 240, 1–10. doi:10.1016/j.epsl.2005.09.038

Środoń, J., 2006. Identification and quantitative analysis of clay minerals, in: Bergaya, F., Theng, B.K.G., Lagaly, G. (Eds.), Developments in Clay Science. pp. 765–787. doi:10.1016/S1572- 4352(05)01028-7

Stack, K.M., Milliken, R.E., 2015. Modeling near-infrared reflectance spectra of clay and sulfate mixtures and implications for Mars. Icarus 250, 332–356. doi:10.1016/j.icarus.2014.12.009

Stahly, G.P., 2012. Advantages of a Cu vs. Co X-ray Diffraction Source, Triclinic Labs. doi:http://www.tricliniclabs.com/downloadable- documents/Advantages%20of%20a%20Cu%20vs.%20Co%20X- ray%20Diffraction%20Source%20-%20Stahly%20-%20Triclinic%20Labs%20-Q32012.pdf

Steudel, R., 1996. Mechanism for the Formation of Elemental Sulfur from Aqueous Sulfide in Chemical and Microbiological Desulfurization Processes. Ind. Eng. Chem. Res. 35, 1417– 1423. doi:10.1021/ie950558t

200

Stoffregen, R.E., 1993. Stability relations of jarosite and natrojarosite at 150-250° C. Geochim. Cosmochim. Acta 57, 2417–2429.

Stoffregen, R.E., Alpers, C.N., 1992. Observations on the unit-cell dimensions, H2O contents, and 6D values of natural and synthetic alunite. Am. Mineral. 77, 1092–1098.

Stumm, W., Morgan, J.J., 1996. Aquatic Chemistry: Chemical Equilibria and Rates in Natural Waters, 3rd ed. Wiley-Interscience, Hoboken, NJ.

Stuurman, C.M., Osinski, G.R., Holt, J.W., Levy, J.S., Brothers, T.C., Kerrigan, M., Campbell, B.A., 2016. SHARAD detection and characterization of subsurface water ice deposits in Utopia Planitia, Mars. Geophys. Res. Lett. 43. doi:10.1002/2016GL070138.Received

Summons, R.E., Amend, J.P., Bish, D., Buick, R., Cody, G.D., Des Marais, D.J., Dromart, G., Eigenbrode, J.L., Knoll, A.H., Sumner, D.Y., 2011. Preservation of Martian Organic and Environmental Records: Final Report of the Mars Biosignature Working Group. Astrobiology 11, 157–181. doi:10.1089/ast.2010.0506

Swayze, G.A., Clark, R.N., Goetz, A.F.H., Livio, K.E., Breit, G.N., Kruse, F.A., Sutley, S.J., Snee, L.W., Lowers, H.A., Post, J.L., Stoffregen, R.E., Ashley, R.P., 2014. Mapping Advanced Argillic Alteration at Cuprite, Nevada, Using Imaging Spectroscopy. Econ. Geol. 109, 1179– 1221.

Swayze, G.A., Clark, R.N., Sutley, S.J., Gent, C.A., Rockwell, B.W., Blaney, D.L., Post, J.L., Farm, B.P., 2002. Mineral Mapping Mauna Kea and Mauna Loa Shield Volcanos on Hawaii Using AVIRIS Data and the USGS Tetracorder Spectral Identification System: Lessons Applicable to the Search for Relict Martian Hydrothermal Systems, in: Proceedings of the 11th JPL Airborne Earth Science Workshop. pp. 373–387.

Swayze, G.A., Ehlmann, B.L., Milliken, R.E., Poulet, F., Wray, J.J., Rye, R.O., Clark, R.N., Desborough, G.A., Crowley, J.K., Gondet, B., Mustard, J.F., Seelos, K.D., Murchie, S.L., 2008. Discovery of the Acid-Sulfate Mineral Alunite in Terra Sirenum, Mars, Using MRO CRISM: Possible Evidence for Acid-Saline Lacustrine Deposits?, in: American Geophysical Union, Fall Meeting. p. abstract #P44A-04.

Swayze, G.A., Smith, K.S., Clark, R.N., Sutley, S.J., Pearson, R.M., Vance, J.S., Hageman, P.L., Briggs, P.H., Meier, A.L., Singleton, M.J., Roth, S., 2000. Using Imaging Spectroscopy to Map Acidic Mine Waste. Environ. Sci. Technol. 34, 47–54.

Szwast, M.A., Richardson, M.I., Vasavada, A.R., 2006. Surface dust redistribution on Mars as observed by the Mars Global Surveyor and Viking orbiters. J. Geophys. Res. 111. doi:10.1029/2005JE002485

Tanaka, B.K.L., Skinner, J.A., Hare, T.M., 2005. Geologic Map of the Northern Plains of Mars.

Taylor, G.J., 2013. The bulk composition of Mars. Chemie der Erde 73, 401–420. doi:10.1016/j.chemer.2013.09.006

201

Thollot, P., Mangold, N., Ansan, V., Le Mouélic, S., Milliken, R.E., Bishop, J.L., Weitz, C.M., Roach, L.H., Mustard, J.F., Murchie, S.L., 2012. Most Mars minerals in a nutshell: Various alteration phases formed in a single environment in Noctis Labyrinthus. J. Geophys. Res. 117. doi:10.1029/2011JE004028

Thollot, P., Mangold, N., LeMouélic, S., Milliken, R.E., Roach, L.H., Mustard, J.F., 2011. A large diversity of hydrated minerals formed from interaction with hydrothermal fluids in Noctis Labyrinthus, Mars, in: EPSC-DPS Joint Meeting. p. EPSC-DPS2011-1067.

Thordarson, T., Larsen, G., 2007. Volcanism in Iceland in historical time: Volcano types, eruption styles and eruptive history. J. Geodyn. 43, 118–152. doi:10.1016/j.jog.2006.09.005

Thorseth, I.H., Furnes, H., Tumyr, O., 1991. A textural and chemical study of Icelandic palagonite of varied composition and its bearing on the mechanism of the glass-palagonite transformation. Geochim. Cosmochim. Acta 55, 731–749.

Torrent, J., Guzman, R., Parra, M.A., 1982. Influence of Relative Humidity on the Crystallization of Fe (III) Oxides From Ferrihydrite. Clays Clay Miner. 30, 337–340. doi:10.1346/CCMN.1982.0300503

Tosca, N.J., McLennan, S.M., Dyar, M.D., Sklute, E.C., Michel, F.M., 2008. Fe oxidation processes at Meridiani Planum and implications for secondary Fe mineralogy on Mars. J. Geophys. Res. E Planets 113, 1–20. doi:10.1029/2007JE003019

Tosca, N.J., McLennan, S.M., Lindsley, D.H., Schoonen, M.A.A., 2004. Acid-sulfate weathering of synthetic Martian basalt: The acid fog model revisited. J. Geophys. Res. 109. doi:10.1029/2003JE002218

Treiman, A.H., Amundsen, H.E.F., Blake, D.F., Bunch, T., 2002. Hydrothermal origin for carbonate globules in Martian meteorite ALH84001: A terrestrial analogue from Spitsbergen (Norway). Earth Planet. Sci. Lett. 204, 323–332. doi:10.1016/S0012-821X(02)00998-6

Treiman, A.H., Bish, D.L., Vaniman, D.T., Chipera, S.J., Blake, D.F., Ming, D.W., Morris, R. V, Bristow, T.F., Morrison, S.M., Baker, M.B., Rampe, E.B., Downs, R.T., Filiberto, J., Glazner, A.F., Gellert, R., Thompson, L.M., Schmidt, M.E., Le Deit, L., Wiens, R.C., McAdam, A.C., Achilles, C.N., Edgett, K.S., Farmer, J.D., Fendrich, K. V., Grotzinger, J.P., Gupta, S., Morookian, J.M., Newcombe, M.E., Rice, M.S., Spray, J.G., Stolper, E.M., Sumner, D.Y., Vasavada, A.R., Yen, A.S., 2016. Mineralogy, provenance, and diagenesis of a potassic basaltic sandstone on Mars: CheMin X-ray diffraction of the Windjana sample (Kimberley area, Gale Crater). J. Geophys. Res. Planets 121, 75–106. doi:10.1002/2015JE004932

Treiman, A.H., Robinson, K.L., Blake, D.F., Bish, D., 2010. Mineralogy Determinations by CheMin XRD, Tested on Ultramafic Rocks (Mantle Xenoliths), in: Astrobiology Science Conference. pp. 5351–5352. doi:10.1016/j.orggeochem.2009.08.003.

Trønnes, R.G., 1990. Basaltic Melt Evolution of the Hengill Volcanic System, SW Iceland, and Evidence for Clinopyroxene Assimilation in Primitive Tholeiitic Magmas. J. Geophys. Res. 95, 15893–15910. doi:10.1029/JB095iB10p15893

202

Trukhin, V.I., Zhilyayeva, V.A., Safroshkin, V.Y., Genshaft, Y.S., Gorshkov, A.G., 1995. Highly ferrous basalts of Iceland and their magnetic properties. Phys. Solid Earth 31, 24–37.

Vandenabeele, P., 2013. Practical Raman Spectroscopy – An Introduction. John Wiley & Sons, Inc.

Vaniman, D.T., Bish, D.L., Ming, D.W., Bristow, T.F., Morris, R. V., Blake, D.F., Chipera, S.J., Morrison, S.M., Treiman, A.H., Rampe, E.B., Rice, M., Achilles, C.N., Grotzinger, J.P., McLennan, S.M., Williams, J., Bell, J.F., Newsom, H.E., Downs, R.T., Maurice, S., Sarrazin, P., Yen, A.S., Morookian, J.M., Farmer, J.D., Stack, K., Milliken, R.E., Ehlmann, B.L., Sumner, D.Y., Berger, G., Crisp, J.A., Hurowitz, J.A., Anderson, R., Des Marais, D.J., Stolper, E.M., Edgett, K.S., Gupta, S., Spanovich, N., the MSL Science Team, 2014. Mineralogy of a Mudstone at Yellowknife Bay, Gale Crater, Mars. Science (80-. ). 343. doi:10.1126/science.1243480

Walker, J.A., Carr, M.J., 1986. Compositional variations caused by phenocryst sorting at Cerro Negro volcano, Nicaragua. Geol. Soc. Am. Bull. 97, 1156–1162. doi:10.1130/0016- 7606(1986)97<1156:CVCBPS>2.0.CO;2

Walker, J.J., Spear, J.R., Pace, N.R., 2005. Geobiology of a microbial endolithic community in the Yellowstone geothermal environment. Nature 434, 1011–1014. doi:10.1038/nature03499.1.

Wang, A., Freeman, J.J., Jolliff, B.L., Chou, I.-M., 2006a. Sulfates on Mars: A systematic Raman spectroscopic study of hydration states of magnesium sulfates. Geochim. Cosmochim. Acta 70, 6118–6135. doi:10.1016/j.gca.2006.05.022

Wang, A., Haskin, L.A., Squyres, S.W., Jolliff, B.L., Crumpler, L., Gellert, R., Schröder, C., Herkenhoff, K., Hurowitz, J., Tosca, N.J., Farrand, W.H., Anderson, R., Knudson, A.T., 2006b. Sulfate deposition in subsurface regolith in Gusev crater, Mars. J. Geophys. Res. E Planets 111, 1–19. doi:10.1029/2005JE002513

Wang, A., Jolliff, B.L., Haskin, L.A., 1999. Raman spectroscopic characterization of a highly weathered basalt: Igneous mineralogy, alteration products, and a microorganism. J. Geophys. Res. 104, 27067–27077. doi:10.1029/1999JE001071

Wang, A., Korotev, R.L., Jolliff, B.L., Haskin, L.A., Crumpler, L., Farrand, W.H., Herkenhoff, K.E., de Souza, P., Kusack, A.G., Hurowitz, J.A., Tosca, N.J., 2006c. Evidence of phyllosilicates in Wooly Patch, an altered rock encountered at West Spur, Columbia Hills, by the Spirit rover in Gusev crater, Mars. J. Geophys. Res. 111. doi:10.1029/2005JE002516

Wanke, H., Bruckner, J., Dreibus, G., Rieder, R., Ryabchikov, I., 2001. Chemical composition of rocks and soils at the Pathfinder site. Space Sci. Rev. 96, 317–330. doi:10.1023/A

Warner, N.H., Farmer, J.D., 2010. Subglacial hydrothermal alteration minerals in Jökulhlaup deposits of Southern Iceland, with implications for detecting past or present habitable environments on Mars. Astrobiology 10, 523–547. doi:10.1089/ast.2009.0425

Water, M.R., Des Marais, D.J., 1993. Preservation of Biological Information in Thermal Spring

203

Deposits: Developing a Strategy for the Search for Fossil Life on Mars. Icarus.

Watters, T.R., Maxwell, T.A., 1986. Orientation, Relative Age, and Extent of the Tharsis Plateau Ridge System. J. Geophys. Res. 91, 8113–8125.

Wdowiak, T.J., Klingelhöfer, G., Wade, M.L., Nunez, J.I., 2003. Extracting science from Mössbauer spectroscopy on Mars. J. Geophys. Res. 108. doi:10.1029/2003JE002071

Weitz, C.M., Bishop, J.L., Baker, L.L., Berman, D.C., 2014. Fresh exposures of hydrous Fe- bearing amorphous silicates on Mars. Geophys. Res. Lett. 41, 8744–8751. doi:10.1002/2014GL062065

Weitz, C.M., Bishop, J.L., Grant, J. a., 2013. Gypsum, opal, and fluvial channels within a trough of noctis labyrinthus, mars: Implications for aqueous activity during the late hesperian to amazonian. Planet. Space Sci. 87, 130–145. doi:10.1016/j.pss.2013.08.007

Weitz, C.M., Bishop, J.L., Roach, L., Milliken, R.E., Rodriguez, J.A., 2010. Mineralogy and Morphology of Light-Tones Deposits in Noctis Labyrinthus, in: Lunar and Planetary Science Conference. doi:10.1016/geomorph.2007.04.017.

Weitz, C.M., Bishop, J.L., Thollot, P., Mangold, N., Roach, L.H., 2011. Diverse mineralogies in two troughs of Noctis Labyrinthus, Mars. Geology 39, 899–902. doi:10.1130/G32045.1

Weitz, C.M., Lane, M.D., Staid, M., Dobrea, E.N., 2008. Gray hematite distribution and formation in Ophir and Candor chasmata. J. Geophys. Res. E Planets 113, 1–30. doi:10.1029/2007JE002930

Werner, S.C., 2009. The global martian volcanic evolutionary history. Icarus 201, 44–68. doi:10.1016/j.icarus.2008.12.019

Wiens, R.C., Maurice, S., McCabe, K., Cais, P., Anderson, R.B., Beyssac, O., Bonal, L., Clegg, S., Deflores, L., Dromart, G., Fischer, W.W., Forni, O., Gasnault, O., Grotzinger, J.P., Johnson, J.R., Martinez-Frias, J., Mangold, N., McLennan, S., Montmessin, F., Rull, F., Sharma, S.K., Sautter, V., Lewin, E., Cloutis, E., Poulet, F., Bernard, S., McConnochie, T., Lanza, N., Newsom, H., Ollila, A., Leveille, R., Le Mouelic, S., Lasue, J., Melikechi, N., Meslin, P.-Y., Misra, A., Grasset, O., Angel, S.M., Fouchet, T., Beck, P., Bridges, N., Bousquet, B., Fabre, C., Pinet, P., Benzerara, K., Montagnac, G., 2016. The SuperCam Remote Sensing Instrument Suitefor Mars 2020, in: Lunar and Planetary Science Conference. pp. 1322–1323.

Williams, R.M.E., Grotzinger, J.P., Dietrich, W.E., Gupta, S., Sumner, D.Y., Wiens, R.C., Mangold, N., Malin, M.C., Edgett, K.S., Maurice, S., Forni, O., Gasnault, O., Ollila, A., Newsom, H.E., Dromart, G., Palucis, M.C., Yingst, R.A., Anderson, R.B., Herkenhoff, K.E., Le Mouélic, S., Goetz, W., Madsen, M.B., Koefoed, A., Jensen, J.K., Bridges, J.C., Schwenzer, S.P., Lewis, K.W., Stack, K.M., Rubin, D., Kah, L.C., Bell, J.F., Farmer, J.D., Sullivan, R., Van Beek, T., Blaney, D.L., Pariser, O., Deen, R.G., MSL Science Team, 2013. Martian Fluvial Conglomerates at Gale Crater. Science (80-. ). 340, 1068–1072. doi:10.1126/science.1237317

204

Wilson, L., Ghatan, G.J., Head, J.W., Mitchell, K.L., 2004. Mars outflow channels: A reappraisal of the estimation of water flow velocities from water depths, regional slopes, and channel floor properties. J. Geophys. Res. 109. doi:10.1029/2004JE002281

Wilson, L., Head, J.W., 1994. Mars: Review and analysis of volcanic eruption theory and relationships to observed landforms. Rev. Geophys. 32, 221–263.

Wilson, S.A., Howard, A.D., Moore, J.M., Grant, J.A., 2016. A cold-wet middle latitude environment on Mars during the Hesperian-Amazonian transition: Evidence from northern Arabia valleys and paleolakes. J. Geophys. Res. Planets 121. doi:10.1002/2016JE005052.Received

Witze, A., 2017. NASA narrows sites for rock-harvesting Mars rover. Nature 542, 279–280.

Wolfe, C.J., Bjarnason, I.T., Vandecar, J.C., Solomon, S.C., 2002. Assessing the depth resolution of tomographic models of upper mantle structure beneath Iceland. Geophys. Res. Lett. 29. doi:10.1029/2001GL013657

Wray, J.J., Hansen, S.T., Dufek, J., Swayze, G.A., Murchie, S.L., Seelos, F.P., Skok, J.R., Irwin, R.P., Ghiorso, M.S., 2013. Prolonged magmatic activity on Mars inferred from the detection of felsic rocks. Nat. Geosci. 6, 1013–1017. doi:10.1038/ngeo1994

Wray, J.J., Milliken, R.E., Murchie, S.L., Swayze, G.A., Dundas, C.M., Seelos, F.P., Squyres, S.W., 2008. Clays and Sulfates in a Potential Lacustrine Evaporite Sequence at Columbus Crater, Mars, in: American Geophysical Union, Fall Meeting. p. P53B–1446.

Wray, J.J., Murchie, S.L., Bishop, J.L., Ehlmann, B.L., Milliken, R.E., Wilhelm, M.B., Seelos, K.D., Chojnacki, M., 2016. Orbital evidence for more widespread carbonate-bearing rocks on Mars. J. Geophys. Res. Planets. doi:10.1002/2015JE004972

Wray, J.J., Murchie, S.L., Squyres, S.W., Seelos, F.P., Tornabene, L.L., 2009. Diverse aqueous environments on ancient Mars revealed in the southern highlands. Geology 37, 1043–1046. doi:10.1130/G30331A.1

Wray, J.J., Noe Dobrea, E.Z., Arvidson, R.E., Wiseman, S.M., Squyres, S.W., McEwen, a. S., Mustard, J.F., Murchie, S.L., 2009. Phyllosilicates and sulfates at Endeavour Crater, Meridiani Planum, Mars. Geophys. Res. Lett. 36. doi:10.1029/2009GL040734

Wyatt, M.B., McSween, H.Y., Tanaka, K.L., Head, J.W., 2004. Global geologic context for rock types and surface alteration on Mars. Geology 32, 645–648. doi:10.1130/G20527.1

Yant, M., Rogers, A.D., Nekvasil, H., Zhao, Y.-Y.S., Bristow, T., 2016. Spectral characterization of acid weathering products on Martian basaltic glass. J. Geophys. Res. Planets 121, 516– 541. doi:10.1002/2014JD022076.Received

Yen, A.S., Morris, R. V., Clark, B.C., Gellert, R., Knudson, A.T., Squyres, S., Mittlefehldt, D.W., Ming, D.W., Arvidson, R., McCoy, T., Schmidt, M., Hurowitz, J., Li, R., Johnson, J.R., 2008. Hydrothermal processes at Gusev Crater: An evaluation of Paso Robles class soils. J.

205

Geophys. Res. 113. doi:10.1029/2007JE002978

206