is clear that subsurface water ice may be extensive and may harbour liquid water ...... Bish, D. L., J. William Carey, D. T. Vaniman, and S. J. Chipera (2003), ...
Two complementary approaches in refining the search for liquid water and habitable environments on present-day Mars. Eriita G. Jones A thesis submitted for the degree of Doctor of Philosophy Of The Australian National University
Research School of Astronomy and Astrophysics March 2012
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Declaration I hereby declare that the work in this thesis undertaken in the period of 2008 to 2012 is that of the candidate alone, except where indicated below. Much of Chapter 3 and Chapter 4 has been published as a paper in Astrobiology. The paper was co-authored by Charles Lineweaver. Additional results from Chapter 4 have been published as a paper in the Australian Journal of Earth Sciences. The paper was co-authored by Charles Lineweaver. Some of the background material in Chapter 2 (2.4) and Chapter 5 (5.1) has been published as a paper in the Astronomical Society of the Pacific Conference Series. The paper was coauthored by Charles Lineweaver. Much of Chapter 6 has been published as a paper in Astrobiology. It was co-authored by Charles Lineweaver and Jonathan Clarke. Some of the background material in Chapter 7 (7.1.2, 7.1.3) and methodology in Chapter 8 (8.2) has been published as a paper in the Australian Space Sciences Conference proceedings. The paper was co-authored by Franklin Mills, Graziella Caprarelli, Bruce Doran and Jonathan Clarke. Some of the method and preliminary work from Chapter 8 has been published as an abstract in the Lunar and Planetary Science Conference program. This was co-authored by Franklin Mills, Bruce Doran, Graziella Caprarelli and Jonathan Clarke. Much of the results in Chapter 8 have been written in a manuscript that is currently in preparation. This is co-authored by Franklin Mills, Bruce Doran, Graziella Caprarelli and Jonathan Clarke. In all occasions, co-authors provided suggestions and comments for the manuscripts and helped in the interpretation of the results.
__________________ Eriita Jones
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Abstract All known active life requires liquid water. The correlation between liquid water and the presence of life on Earth has guided the search for life on other planets. For terrestrial-like life to exist in the harsh conditions that dominate the surfaces of other rocky planets, the minimum fundamental requirements of liquid water, nutrients, and chemical energy must be met. Within our solar system, Mars is a strong candidate for hospitable environments able to support life, due to the reservoirs of water within its crust and the strong likelihood of liquid water. The aim of this thesis is to refine the search for liquid water and environments that may be hospitable to life on Mars. Two complementary methodologies are developed and utilised to achieve this aim. Water requires a relatively narrow range of pressures and temperatures to occur in the liquid phase. The first approach of this thesis compares this range with the pressure-temperature conditions that occur within the Earth, the Earth’s active biosphere, and Mars. Temperature, pressure and water activity are examined to determine the extent to which they restrict life from some liquid water environments. The relevant thresholds are then applied to Mars and compared to models of where liquid water environments are likely to occur under present-day martian conditions. Extensive regions of the Earth may be inhospitable despite lying within the hydrosphere. Life is likely restricted from ~ 81% of the volume of the hydrosphere of Earth due to high temperature and/or low water activity. In contrast, the fraction of Mars that can support liquid water is five times larger than that of Earth, given estimates of an average martian brine. Many environments within the martian crust can potentially support life, with perennially habitable conditions extending from approximately 10 to 37 km beneath the surface. The surface and shallow regolith may also be habitable in the warmest regions of the planet. The second approach focuses on the shallow subsurface of Mars within the top ~20 m. The thermal behaviour of surface materials determines the occurrence of transient shallow liquid water and habitable temperatures for life. Ten classes of surface materials are identified from analysis of global martian thermal inertia and albedo, through the technique of algorithmic classification. These classes are interpreted as mixtures of dust, sand, duricrust, rocks and ice on the surface, and validated through comparisons with independent datasets. Low latitude locations of dark sand, duricrust and pebbles in Syrtis Major, Oxia Palus, Mawrth Vallis and eastern Meridiani Planum are identified as having high potential for hospitable liquid water environments at < 10 m depth. Dark, coarse, sand dominated surfaces are found in Syrtis Major and Aram Chaos and are predicted to be locations of low volume flows of liquid water, potentially analogous to the observed martian recurring slope lineae. This thesis identifies where habitable liquid water environments may occur on Mars, strengthening the astrobiological significance of the planet and providing direction for future robotic and satellite missions searching for life.
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‘Somewhere in a private place, she packs her bag for outer space. And now she’s waiting for the right kind of pilot to come. She’s saying, ‘I will fly you to the moon and back if you’ll be, if you’ll be my baby. I’ve got a ticket for a world where we belong, if you’ll be my baby’. Her friends are saying that she’s hanging all her hopes on the stars.’ - Darren Hayes, ‘To The Moon and Back’
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Table of Contents 1.
Introduction .................................................................................................................................. 1 1.1.
Approach 1 - assessing global habitability ............................................................... 5
1.2.
Approach 2 – constraining near-surface habitability .......................................... 7
1.3.
Summary ............................................................................................................................... 8
1.4.
Overview & structure ....................................................................................................... 9
1.4.1. Chapter 2- Literature Review & Background: Quantifying Habitability & the Extreme Limits of Life ................................................................................................... 9 1.4.2.
Chapter 3- Phase Diagram of Earth & Water ................................................ 10
1.4.3.
Chapter 4- Phase Diagram of Earth’s Biosphere ......................................... 10
1.4.4.
Chapter 5- Literature Review & Background: Liquid Water on Mars. 11
1.4.5. Chapter 6- Phase Diagram of Mars, Water & the Potential Martian Biosphere ..................................................................................................................................... 11 1.4.6. Chapter 7- Bridging Chapter: The Influence of Surface Materials on the Subsurface Environment ................................................................................................ 12 1.4.7. Chapter 8- Remote Sensing of Thermal Materials on the Martian Surface 12 1.4.8.
Chapter 9- Synthesis & Conclusions................................................................. 13
1.5.
Publication summary...................................................................................................... 13
1.6.
Reference material .......................................................................................................... 14
2. Literature Review & Background: Quantifying Habitability & the Extreme Limits of Life ........................................................................................................................................ 21 2.1.
First order: habitable zone around a star ............................................................... 22
2.2.
Fluids for life ...................................................................................................................... 25
2.3.
Energy & nutrients .......................................................................................................... 26
2.4.
Phase requirements of liquid water ......................................................................... 28
2.4.1.
Interpreting the water phase diagram ............................................................ 29
2.4.2.
Freezing point depression and boiling point elevation ............................ 30
2.5.
Water activity .................................................................................................................... 33
2.5.1.
Water activity of solutions (solute effect) ..................................................... 35
2.5.2.
Thin films (temperature effect) ......................................................................... 36
2.5.3.
Implications for life................................................................................................. 38
2.6.
Pressure ............................................................................................................................... 41
2.6.1.
High threshold .......................................................................................................... 41
2.6.2.
Low threshold ........................................................................................................... 42
2.7.
Temperature ...................................................................................................................... 43
2.7.1.
High threshold .......................................................................................................... 43
2.7.2. 2.8.
Other factors ...................................................................................................................... 45
2.8.1. 2.9. 3.
Low threshold .......................................................................................................... 44 Size of life ................................................................................................................... 46
Summary ............................................................................................................................. 47
Phase Diagram of Earth & Water........................................................................................ 49 3.1.
Development of the model ........................................................................................... 51
3.1.1.
Ocean water & largest liquid phase space ..................................................... 51
3.1.2.
Earth phase model .................................................................................................. 52
3.1.2.a. 3.1.3.
Atmosphere & oceans ........................................................................................... 54
3.1.4.
Surface & crust ......................................................................................................... 55
3.1.4.a.
Boundary ........................................................................................................... 55
3.1.4.b.
Average .............................................................................................................. 57
3.1.4.c.
Geothermal gradients in the crust ........................................................... 59
3.1.5.
Mantle & core ........................................................................................................... 61
3.1.6.
Transient .................................................................................................................... 64
3.1.7.
Interpreting the Earth phase diagram ............................................................ 64
3.2.
3.1.7.a.
Temporal variation ........................................................................................ 64
3.1.7.b.
The meaning of pressure ............................................................................. 65
Results: the pressure-temperature representation of Earth & liquid water 66
3.2.1.
Implications of high solute & thin film liquids ............................................ 67
3.2.2.
Deepest liquid water.............................................................................................. 70
3.2.2.a. 3.2.3. 3.3. 4.
Pressure, temperature & depth ................................................................ 54
Pore space & fluid access ............................................................................ 72
Liquid water budget............................................................................................... 73
Summary ............................................................................................................................. 74
Phase Diagram of Earth’s Biosphere ................................................................................. 75 4.1.
Development of the biosphere model...................................................................... 77
4.1.1.
Extrapolating to define maximal life space ................................................... 79
4.1.2.
Active & dormant life ............................................................................................. 80
4.1.2.a. 4.2.
Volume calculations ...................................................................................... 83
Results: the pressure temperature representation of the biosphere.......... 84
4.2.1.
Extent of the biosphere......................................................................................... 87
4.2.1.a.
Pore space & permeability .......................................................................... 88
4.2.2.
The subzero biosphere ......................................................................................... 90
4.2.3.
The ambiguity of cold atmospheric life .......................................................... 91
4.2.4.
Uninhabited liquid water & its implications for the search for life ..... 92
4.2.5.
Estimating the amount of uninhabited water .............................................. 94
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4.3. 5.
Summary ............................................................................................................................. 96
Literature Review & Background: Liquid Water on Mars......................................... 99 5.1.
Water stability ................................................................................................................ 100
5.1.1.
Ice ............................................................................................................................... 100
5.1.2.
Liquid ........................................................................................................................ 104
5.2.
Reservoirs of present day water ice ...................................................................... 107
5.2.1.
Polar caps ................................................................................................................ 107
5.2.2.
Surface frost............................................................................................................ 108
5.2.3.
Subsurface water ice ........................................................................................... 110
5.2.4.
Surface morphologies indicating subsurface ice ...................................... 116
5.3.
5.2.4.a.
Rampart craters ........................................................................................... 116
5.2.4.b.
Polygons, pingoes ........................................................................................ 119
Present day liquid water ............................................................................................ 120
5.3.1.
Deliquescence & water adsorption................................................................ 120
5.3.2.
Thin films ................................................................................................................. 122
5.3.3.
Brines ........................................................................................................................ 124
5.3.4.
Surface fluvial morphologies ........................................................................... 125
5.3.4.a.
Gullies ............................................................................................................... 126
5.3.4.b.
Dune flows ...................................................................................................... 133
5.3.4.c.
Recurring slope lineae ............................................................................... 134
5.3.5. 5.4. 6.
Beneath the cryosphere ..................................................................................... 137
Summary .......................................................................................................................... 137
Phase Diagram of Mars, Water & the Potential Martian Biosphere ................... 139 6.1.
Development of model ................................................................................................ 141
6.1.1.
Average brine & largest liquid phase space ............................................... 141
6.1.1.a. 6.1.2.
Stability & subsurface closure ................................................................ 142
Mars model ............................................................................................................. 144
6.1.2.a.
Pressure & depth ......................................................................................... 145
6.1.3.
Atmosphere ............................................................................................................ 145
6.1.4.
Surface & crust ...................................................................................................... 146
6.1.4.a.
Boundary......................................................................................................... 146
6.1.4.b.
Mean & variation.......................................................................................... 148
6.1.4.c.
Geothermal gradients in the crust ........................................................ 150
6.1.5.
Mantle & core ......................................................................................................... 152
6.1.6.
Interpreting the Mars phase diagram .......................................................... 155
6.1.6.a.
Transient environments ........................................................................... 155
6.1.6.b.
The meaning of pressure .......................................................................... 156
6.1.7.
The P-T space of Earth’s life ............................................................................. 157
6.2.
Results: the pressure-temperature representation of Mars & liquid water 158
6.2.1.
Implications of high solute & thin film liquids ..........................................160
6.2.2.
Deepest liquid water............................................................................................163
6.2.2.a.
Pore space & fluid access ..........................................................................164
6.3.
Results: the potential martian biosphere ............................................................. 166
6.4.
Summary ...........................................................................................................................174
7. Bridging Chapter: The Influence of Surface Materials on the Subsurface Environment .....................................................................................................................................175 7.1.
7.1.1.
Subsurface temperature profiles ....................................................................177
7.1.2.
Thermal inertia ......................................................................................................180
7.1.3.
Albedo .......................................................................................................................184
7.1.4.
Groundtruthing ......................................................................................................186
7.1.5.
Key martian surface components ...................................................................188
7.2.
7.1.5.a.
Dust ....................................................................................................................188
7.1.5.b.
Sand ...................................................................................................................189
7.1.5.c.
Duricrust ..........................................................................................................190
7.1.5.d.
Rock ...................................................................................................................193
7.1.5.e.
Ice .......................................................................................................................194
Illustration: Estimating the depth to the melting isotherm ..........................195
7.2.1.
General trends ........................................................................................................195
7.2.2.
Example of Gale Crater .......................................................................................197
7.3. 8.
Background ......................................................................................................................177
Summary ...........................................................................................................................205
Remote Sensing of Thermal Materials on the Martian Surface ............................207 8.1.
Background ......................................................................................................................209
8.1.1.
Previous maps of martian thermophysical units......................................209
8.1.2.
Motivation for a different methodology .......................................................212
8.2.
Methodology ....................................................................................................................214
8.2.1.
Observations ...........................................................................................................214
8.2.1.a.
Thermal inertia & albedo ..........................................................................214
8.2.1.b.
Other data........................................................................................................216
8.2.1.c.
Data processing .............................................................................................218
8.2.2.
Algorithmic classification ..................................................................................218
8.2.2.a.
Success in Earth remote sensing ............................................................219
8.2.2.b.
Unsupervised clustering algorithm: ISODATA .................................223
8.2.2.c.
Supervised classifier: MAXLIKE..............................................................227
8.2.3. Caveats to the application of ISODATA & Maximum Likelihood algorithms .................................................................................................................................229
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8.3.
8.2.3.a.
Number of classes ........................................................................................ 229
8.2.3.b.
Variable weighting ...................................................................................... 232
8.2.3.c.
Data distribution .......................................................................................... 233
8.2.3.d.
Assessing classification accuracy .......................................................... 233
Results: the spatial distribution of surface materials ..................................... 237
8.3.1.
Global trends .......................................................................................................... 243
8.3.1.a. Comparison between thermophysical classes & major geological structures 244 8.3.1.b. 8.3.2.
9.
Interpretation of thermophysical classes ................................................... 248
8.3.2.a.
Classes 1, 2 & 3: sand.................................................................................. 249
8.3.2.b.
Classes 1, 3 & 6: pebbles, duricrust. ..................................................... 253
8.3.2.c.
Classes 4-7: fine-medium sand & pebbles, varying dust .............. 255
8.3.2.d.
Class 5: possible subsurface ice ............................................................. 257
8.3.2.e.
Class 8: dust ................................................................................................... 259
8.3.2.f.
Classes 9 & 10: ice, polar caps & bedrock .......................................... 260
8.3.3. 8.4.
Comparison between thermophysical classes & geologic age ... 247
Comparison to previous work ......................................................................... 262
Summary .......................................................................................................................... 265
Synthesis & Conclusions...................................................................................................... 267 9.1.
Overview of results....................................................................................................... 269
9.1.1.
Earth’s hydrosphere (Chapter 3) ................................................................... 269
9.1.2.
The limits of life & the extent of Earth’s biosphere (Chapter 4) ........ 270
9.1.3.
Mars’ potential hydrosphere (Chapter 6, 7) .............................................. 271
9.1.4.
Potential biosphere of Mars (Chapter 6)..................................................... 272
9.1.5.
Distribution of martian surface materials (Chapter 8) .......................... 273
9.2. Thesis synthesis: potentially habitable environments near the surface of Mars. 274 9.2.1.
Surface & shallow soil ......................................................................................... 275
9.2.1.a.
Liquid water at the surface ...................................................................... 277
9.2.1.b.
Shallow subsurface liquid water ........................................................... 279
9.2.2.
Organisms with relevant adaptations for Mars ........................................ 280
9.2.2.a.
Surface environments ................................................................................ 280
9.2.2.b.
Subsurface environments ......................................................................... 281
9.3.
Conclusions...................................................................................................................... 282
9.4.
Future directions ........................................................................................................... 285
9.4.1.
Habitability of Mars through time ................................................................. 285
9.4.2.
Further characterisation of thermophysical units .................................. 286
9.4.2.a.
Tracing the origin of fines ........................................................................ 286
9.4.2.b.
Understanding the effect of past water activity on grain size ....287
Appendix A: Image Metadata ......................................................................................................291 Appendix B: Processing Procedure for HiRISE & THEMIS Images ..............................297 HiRISE images ..............................................................................................................................297 Processing .................................................................................................................................297 THEMIS images............................................................................................................................298 Processing .................................................................................................................................298 Appendix C: Methodological Considerations for Chapter 8............................................301 Optimal number of thermophysical units .........................................................................301 Alternate classification algorithms ......................................................................................303 Glossary...............................................................................................................................................311 References..........................................................................................................................................316
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List of Figures Figure 1-1: Martian sunset captured by the Mars Exploration Rover Spirit at Gusev (175.5 ˚E, 14.6 ˚S) ....................................................................................................................... 2 Figure 1-2: Two approaches implemented in this thesis ..................................................... 4 Figure 1-3: Pressures and temperatures within the universe .......................................... 6 Figure 1-4: Map of martian surface topography from the Mars Orbiter Laser Altimeter [Smith et al., 2001] .............................................................................................. 14 Figure 1-5: Map of martian thermal inertia derived from Thermal Emission Spectrometer observations [Putzig and Mellon, 2007a] .......................................... 14 Figure 1-6: Map of martian albedo from the Thermal Emission Spectrometer [Putzig and Mellon, 2007a] .................................................................................................. 14 Figure 1-7: Map of major martian geologic units, based on surface age and morphology [Skinner et al., 2006] (details of units given in reference)............. 14 Figure 1-8: Recent obliquity values for Mars from numerical modelling of the martian orbit [Morgan et al., 2011] (Fig. 20) ............................................................... 20 Figure 2-1: Extrasolar planets around solar-mass stars [Schneider et al., 2011] .... 23 Figure 2-2: Exoplanets which may enter their habitable zones, defined by the potential for liquid water on their surface [Kane and Gelino, 2011, 2012] ...... 24 Figure 2-3: Partial pressures of water on Earth .................................................................... 30 Figure 2-4: Schematic of the increase in water activity with the addition of water to a dry material and the thickness of the liquid film [Seiler, 1976] ......................... 35 Figure 2-5: Liquid water content of soil (am) determined by water activity (aw) for temperatures of 200 K (upper curve), 220 K, 240 K and 260 K [Mohlmann, 2008] (Fig. 8) ............................................................................................................................. 38 Figure 3-1: Average surface pressures and temperatures of some solar system planets and moons superimposed on the phase diagram of water ...................... 51 Figure 3-2: Geothermal heat flow map of north America .................................................. 60 Figure 3-3: Phase diagram of water chemistries found on Earth ................................... 67 Figure 3-4: Earth phase diagram superimposed on that of liquid water, with average geotherm and error bars at boundary regions. The second geotherm between ~ 4-10 km ocean depth corresponds to the geotherm through the oceanic crust .............................................................................................................................. 68 Figure 3-5: Simplified plot of Earth’s phase space and the phase space of Earth water ............................................................................................................................................. 70 Figure 4-1: Liquid water, Earth and life .................................................................................... 77 Figure 4-2: Schematic of the method for constructing the biosphere phase model 78 Figure 4-3: Estimated phase space of Earth life superimposed on the P-T diagrams of Earth water from Figure 3-3 ........................................................................................... 84 Figure 4-4: Superposition of inhabited environments from Figure 4-3 on the phase space of the Earth and water from Figure 3-5 .............................................................. 85 Figure 4-5: Deep hot life in restricted pore space ................................................................ 88 Figure 4-6: Deep, icy life from Antarctic ice core. Cell activity could be measured after incubation, indicating that cells in situ were dormant and able to maintain vital functions but not reproduce ................................................................... 91 Figure 4-7: Schematic of Earth's hydrosphere and biosphere ......................................... 95
Figure 4-8: Quantifying uninhabited water. Considering the environments on Earth that have liquid water, two ways of visualising what fraction of these environments are inhabited by life are plotted ........................................................... 96 Figure 5-1: Partial pressures of water on Mars...................................................................102 Figure 5-2: Depth of perennial ice table in the current martian climate, from models of mean surface temperature and thermal inertia, assuming zero elevation and water vapour density of 102 kg/m2 [Mellon et al., 2004b] (Fig. 3). .................................................................................................................................................103 Figure 5-3: A location in Terra Sirenum observed by CRISM, near 38.9 ˚S, 195.9 ˚E .......................................................................................................................................................109 Figure 5-4: Water-equivalent hydrogen concentrations within the top 2 m depth of the martian subsurface from orbital neutron flux [Maurice et al., 2011] (Fig. 32) ...............................................................................................................................................112 Figure 5-5: Lobate debris apron in Niger Vallis, an ancient channel feeding into Hellas Planitia at 35.6 ˚S, 268 ˚W. Feature is likely an ancient glacier covered by rocky regolith ....................................................................................................................115 Figure 5-6: Comparison between martian rampart craters Yuty and Wohoo displaying fluidized ejecta (left; 22.2 ˚N, 325.9 ˚E), and lunar crater Timocharis with dry rayed ejecta (right; 26.7 ˚N, 13.7 ˚W) ..........................................................117 Figure 5-7: Spheroids on the leg of the Phoenix Lander, photographed by the robotic arm camera [Rennó et al., 2009] (Fig. 3) ...................................................... 122 Figure 5-8: Gullies in Corozal crater in Terra Cirrenum at 159.4 ˚E, 38.7 ˚S (HiRISE image) ........................................................................................................................................127 Figure 5-9: North polar gullies at 68.5 ˚S, 1.3˚ E (HiRISE image) .................................130 Figure 5-10: Fresh bright deposits in Gasa crater gullies at 129.4 ˚E, 35.7 ˚S (HiRISE image) ........................................................................................................................................133 Figure 5-11: Active dune gullies in the Kaiser dune field at 46.7 ˚S, 20.1 ˚E [Diniega et al., 2010] (HiRISE image) ..............................................................................................135 Figure 5-12: Active recurring slope lineae observed in the 298 km Newton crater at 41.6 ˚S, 202.3 ˚E .................................................................................................................136 Figure 6-1: Mars Global Surveyor gravitational anomaly map [Smith et al., 1999]. Gravitational anomalies are regions where the gravity and topography are not correlated, and can be used to infer subsurface densities and elastic thickness [McKenzie et al., 2002]. .......................................................................................................149 Figure 6-2: Global atmospheric methane concentration on Mars ...............................156 Figure 6-3: Phase diagram of water relevant to martian environments ...................158 Figure 6-4: Mars and liquid water ............................................................................................160 Figure 6-5: The potential martian biosphere .......................................................................170 Figure 6-6: Comparison between the phase space of Earth and Mars .......................172 Figure 7-1: Example subsurface temperature profiles in dust at 59 ˚N [Mellon and Jakosky, 1993] (Fig. 6) .........................................................................................................180 Figure 7-2: Histograms of the global albedo (top) and nighttime thermal inertia (bottom) from TES ................................................................................................................181 Figure 7-3: Scatterplot of 2.592×107 pixels in the global map of median thermal inertia and albedo derived from thermal and visual reflectance data..............186 Figure 7-4: Histogram of pixels within the thermal inertia range of 5-100 tiu and their corresponding albedo values .................................................................................189 Figure 7-5: Histogram of pixels within the thermal inertia range of 150-400 tiu and their corresponding albedo values .................................................................................190 Figure 7-6: Histogram of pixels within the thermal inertia range of 350-600 tiu and their corresponding albedo values .................................................................................191
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Figure 7-7: Histogram of pixels within the thermal inertia range of 1000-3000 tiu and the corresponding albedo values. These values are consistent with pure rock and bedrock, rock with a mixture of dust, and ice. ........................................ 194 Figure 7-8: Estimated maximum range in subsurface temperatures during a martian year at ~ 30˚ absolute latitude ....................................................................... 198 Figure 7-9: Mosaic of Gale Crater in daytime visual imagery from THEMIS, centred at 5.3 ˚S, 137.7 ˚E ................................................................................................................... 199 Figure 7-10: Correlation between thermal inertia and surface temperature at Gale crater from THEMIS nighttime infrared images ....................................................... 201 Figure 7-11: Temperature variability across four seasons within materials in the Gale crater region.................................................................................................................. 203 Figure 7-12: Variation in grain size at Gale Crater from THEMIS thermal inertia measurements [Hobbs et al., 2010] (Fig. 4)................................................................ 204 Figure 8-1: Thermophysical map of Putzig et al. [2005] (Fig. 5) showing the division of the martian surface into 7 thermophysical units, described in Table 8-1. .............................................................................................................................................. 210 Figure 8-2: Illustration of ISODATA clustering in a two-dimensional dataset [Richards, 1986] (Fig. 9.2) ................................................................................................. 222 Figure 8-3: Schematic of the methodology employed in this study ............................ 224 Figure 8-4: Illustration of MAXLIKE classification [Richards, 1995] (Fig. 8.1) ...... 226 Figure 8-5: Comparison between partitioning of the north polar region by MAXLIKE (above) and ISODATA (below). ................................................................... 228 Figure 8-6: Visual assessment of sensitivity to number of classes output from the ISODATA algorithm .............................................................................................................. 231 Figure 8-7: Global distribution of pixel distance values (black solid line) ............... 234 Figure 8-8: Scatterplot of all pixels showing the partitioning of the dataspace into 10 classes (thermal inertia x-axis is linear above; logarithmic below). The discrete values of thermal inertia below 50 tiu are an artefact of the ‘lookup’ model used to derive thermal inertia (Equation 7-8)............................................. 236 Figure 8-9: Map of the martian surface classified into 10 thermophysical units .. 238 Figure 8-10: Classification confidence thresholds based on pixel parameter distances ................................................................................................................................... 238 Figure 8-11: Histograms of martian thermophysical classes 1-8 in thermal inertia (left) and albedo (right) data ............................................................................................ 242 Figure 8-12: Histograms of martian thermophysical classes 9 & 10 in thermal inertia (left) and albedo (right) data ............................................................................. 243 Figure 8-13: Valles Marineris .................................................................................................... 245 Figure 8-14: The summit of Olympus Mons is distinguished in the thermophysical map. ............................................................................................................................................ 246 Figure 8-15: Relationship between terrain age and thermophysical units ............. 248 Figure 8-16: Fraction of global dune area included within each cluster .................. 250 Figure 8-17: Kaiser dune field ................................................................................................... 251 Figure 8-18: Classification of the six martian landing sites ........................................... 253 Figure 8-19: Duricrust or pebbles? THEMIS infrared images are overlain on the geologic map ........................................................................................................................... 254 Figure 8-20: Path of the Opportunity (above) and Spirit (below) rovers ................ 257 Figure 8-21: Lomonosov crater in the thermophysical map (above) and geologic map (below) with THEMIS infrared images superposed ...................................... 259 Figure 8-22: Korolev crater in the thermophysical map (above) and geologic map (below) with THEMIS infrared image overlaid ......................................................... 261
Figure C-0-1: The partitioning of thermal inertia and albedo values into 10 classes from Chapter 8 ........................................................................................................................302 Figure C-0-2: Density of pixels in class 1 of the thermophysical map ........................305 Figure C-0-3: Density of pixels in class 2 ...............................................................................305 Figure C-0-4: Density of pixels in class 3 ...............................................................................306 Figure C-0-5: Density of pixels in class 4 ...............................................................................306 Figure C-0-6: Density of pixels in class 5 ...............................................................................307 Figure C-0-7: Density of pixels in class 6 ...............................................................................307 Figure C-0-8: Density of pixels in class 7 ...............................................................................308 Figure C-0-9: Density of pixels in class 8 ...............................................................................308 Figure C-0-10: Density of pixels in class 9 .............................................................................309 Figure C-0-11: Density of pixels in class 10 ..........................................................................309 Figure C-0-12: Scatterplot of all pixels showing the partitioning of the dataspace into 5 classes (above) and 15 classes (below) ...........................................................310
xxi
List of Tables Table 1-1: Martian geologic epochs. ........................................................................................... 19 Table 2-1: Freezing point depression (˚C) for an ideal solution...................................... 31 Table 2-2: Boiling point elevation (˚C) for an ideal solution............................................. 32 Table 2-3: Properties of common salts. .................................................................................... 32 Table 2-4: Low water activity environments supporting active life.............................. 33 Table 2-5: Eutectic temperatures and water activity of saturated aqueous solutions of some Earth salts. ................................................................................................................. 37 Table 3-1: Average pressures and temperatures of some solar system planets ...... 50 Table 3-2: Model parameters for Earth’s atmosphere. ....................................................... 55 Table 3-3: Extreme temperatures and pressures in Earth’s oceans and crust .......... 56 Table 3-4: Model parameters for Earth’s crust. ..................................................................... 58 Table 3-5: Model parameters for Earth’s core and mantle. .............................................. 63 Table 3-6: Pressure and depths of features of interest ....................................................... 71 Table 4-1: Extreme active life environments .......................................................................... 81 Table 4-2: Dormant life environments ...................................................................................... 82 Table 4-3: Further extreme habitats for active life .............................................................. 82 Table 4-4: Relevant temperature isotherms and their depths from Earth P-T model ......................................................................................................................................................... 85 Table 4-5: Volume estimates for the known biosphere and plausible hydrosphere ......................................................................................................................................................... 92 Table 5-1: Eutectics of saturated aqueous solutions of Mars candidate salts. ....... 123 Table 5-2: Activity and duration of recent flow features. ............................................... 132 Table 6-1: Average pressures and temperatures on Earth and Mars ........................ 141 Table 6-2: Model parameters for Mars’ atmosphere. ....................................................... 146 Table 6-3: Extreme temperatures and pressures on Mars’ surface and crust ........ 147 Table 6-4: Model parameters for Mars’ crust ...................................................................... 150 Table 6-5: Model parameters for Mars’ mantle and core................................................ 153 Table 6-6: Pressure and depths of features of interest. ................................................... 164 Table 6-7: Volume estimates for the potential martian hydrosphere ....................... 165 Table 6-8: Volume estimates for the potential martian biosphere ............................. 169 Table 7-1: Comparison between ground and orbital thermal inertia and albedo 187 Table 7-2: Model thermal inertia values of dust to sand sized particles. ................. 189 Table 7-3: Characteristics of end-member martian surface materials. ..................... 192 Table 7-4: Model parameters for Mars’ subsurface temperatures. ........................... 196 Table 7-5: Length of martian time periods (in seconds) ................................................ 197 Table 7-6: Properties of Gale crater materials from THEMIS-IR band 9 imagery taken across four seasons. ................................................................................................. 202 Table 8-1: Extreme values of thermophysical units from Putzig et al. [2005] (Tab. 1).................................................................................................................................................. 211 Table 8-2: Unambiguous interpretation of surface materials. ...................................... 213 Table 8-3: Global datasets for Mars. ........................................................................................ 217 Table 8-4: Sensitivity of classification to input parameters. ......................................... 226 Table 8-5: Different partitioning of pixels by ISODATA and MAXLIKE. .................... 229 Table 8-6: Approximating classification accuracy by χ2 with 2 degrees of freedom ...................................................................................................................................................... 233
Table 8-7: Attributes of the 10 thermophysical classes (in the form: mean, standard deviation; minimum-maximum). ....................................................................................237 Table 8-8: Interpretation of the 10 martian thermophysical classes. ........................241 Table 8-9:Comparing the partitioning of pixels into classes in this study and in Putzig et al. [2005]. ...............................................................................................................263 Table 9-1: Factors in the selection of astrobiological targets on Mars to search for present-day life.......................................................................................................................276 Table A1: Assorted HiRISE red wavelength images ..........................................................291 Table A2: Assorted THEMIS images. Image ID’s commencing with ‘I’ indicates infrared wavelength, ‘V’ indicates visual wavelength. ............................................292 Table A3: THEMIS infrared images of Gale crater and surrounding region ............292
1
1. INTRODUCTION
‘I met a traveler from an antique land Who said: Two vast and trunkless legs of stone Stand in the desert. Near them, on the sand, Half sunk, a shattered visage lies, whose frown, And wrinkled lip, and sneer of cold command, Tell that its sculptor well those passions read Which yet survive, stamped on these lifeless things, The hand that mocked them, and the heart that fed; And on the pedestal these words appear: “My name is Ozymandias, king of kings: Look on my works, ye Mighty, and despair!” Nothing beside remains. Round the decay Of that colossal wreck, boundless and bare The lone and level sands stretch far away.’ - Percy Bysshe Shelley, ‘Ozymandias’
2
Introduction
The search for life elsewhere in the universe has enthralled humanity for hundreds, possibly thousands of years. The concept of ‘extraterrestrial life’ has captured audiences through science fiction and motivated scientists.
Space
agencies focused in the early 1960’s on developing missions aiming at detecting life beyond Earth and thereby demonstrating that we are not alone [NASA, 1963]. With the tremendous growth in knowledge of the surfaces and interiors of planets within our solar system, as well as the extraordinary resilience of life within Earth, the search for life continues and remains a primary scientific priority.
In
particular, the past and present habitability of Mars are two of the key questions to be addressed by space missions throughout the next decade [Squyres, 2011].
Figure 1-1: Martian sunset captured by the Mars Exploration Rover Spirit at Gusev (175.5 ˚E, 14.6 ˚S). Image courtesy of NASA/JPL-Caltech. Studies of life on Earth have shown that there is strong potential for any world that harbours liquid water to also harbour life. All known life requires frequent liquid water during its active life cycle [Mottl et al., 2007; Rothschild and Mancinelli, 2001]. Hence the correlation between the presence of liquid water and life is a logical tool in the search for life on other planets. Although other liquids, such as ammonia, have been suggested as a plausible basis for extraterrestrial life [Bains, 2004], liquid water is likely the most abundant liquid within rocky planets [Baross et al., 2007]. Hence NASA and other space agencies have adopted a ‘follow the water’ approach in their search for life elsewhere in the solar system and on potentially habitable exoplanets [Hubbard et al., 2002]. Identifying water in the
3
liquid phase on other solar system planets – in the form of pure liquid, brine, or thin films – provides a valuable constraint on the potential of those planets to support life. Past research on the correlation between liquid water and life on Earth is explained in more detail in Chapter 2. Present-day Mars is the focus of this thesis. Not only is the martian surface environment – being dominated by basaltic rock, with gravity ~
ଵ ଷ
Earth, axial tilt
25˚, temperatures -200 ˚C to 0 ˚C and CO2 dominated atmosphere ~
ଵ
ଵ
Earth
pressure [Carr, 2006] – the most akin to Earth’s within our local neighbourhood, it also records the occurrence of extensive liquid water erosion throughout the past and continuing to the present day [Jakosky, 2007]. Throughout the last ~ 4.5 Ga of martian history, liquid water has left its mark on the surface of the planet – scarring canyons, outflow channels, seas, gullies and flooding craters [Berman and Hartmann, 2002; Frey, 2006]; and leaving extensive deposits of phyllosilicates, haematite, and sulphates through its interaction with martian rock [Gendrin et al., 2005; Hamilton et al., 2008]. However, the conditions that have dominated the surface of Mars for the past 500 million years (Ma) to the present-day result in any liquid water environments being transient [Haberle et al., 2001; Richardson and Mischna, 2005]. Rather than the low subsurface temperatures – which restrict the possible liquid water chemistries to those with subzero eutectic temperatures – it is the low partial pressure of water vapour within the atmosphere which provides the strongest constraint, resulting in an evaporation rate of ~10 times that of Earth at zero elevation/sea level [Sears and Moore, 2005].
Hence, any liquid water
exposed to the martian atmosphere will be unstable and difficult to observe. Despite this transient nature however, evidence is growing that near-surface liquid water is able to accumulate seasonally in sufficient volume to influence surface features [Kereszturi et al., 2011; Kossacki and Markiewicz, 2010; McEwen et al., 2011b]. Furthermore, as we are afforded glimpses beneath the martian surface, it is clear that subsurface water ice may be extensive and may harbour liquid water beneath [Farrell et al., 2009; Mangold, 2012]. Thus, as will be discussed in more detail in Chapter 5, there are indicators from past research that Mars may currently support habitable environments.
4
Introduction
Motivated by the search for life and driven by the potential for hospitable liquid water environments on present-day Mars, the research presented in this thesis addresses the following questions: i.
What is an effective methodology for assessing the potential of a planet to support Earth-like life?
ii.
What is the extent of environments on present-day Mars that are potentially hospitable for life?
iii.
Where on Mars do the best candidate habitable environments presently occur?
iv.
What locations should be targeted by the next decade of robotic explorers searching for existing life on the surface and in the shallow subsurface?
To address these questions, two approaches were employed in this thesis. These are outlined below and in Figure 1-2.
Figure 1-2: Two approaches implemented in this thesis. Near-surface environments within the potential martian biosphere identified through the first approach (1.1) provide the motivation for the second approach (1.2). The two approaches are linked, as the second enables a more detailed assessment of potential hospitable conditions within martian surface materials identified by the first approach.
5
1.1.
Approach 1 - assessing global habitability
In order to assess the habitability of other planets, the minimum criteria for life must first be established through examining the physical attributes and chemical resources which Earth biology requires from its environment. This helps identify the necessary conditions for life – the minimum requirements which must be met on another planet in order for it to support life, given our understanding of the thresholds of organisms on the Earth. This approach is inevitably Earth-centred, but that is an unavoidable consequence of only having one sample (albeit a large one) on which to base our theories. Hence the strategy of ‘follow the liquid water’ can be refined by including other fundamental requirements of life. Apart from liquid water, life must also have access to a minimum quota of nutrients to build biomass [Schuur, 2011] and a gradient in chemical energy to fuel metabolism [Hoehler, 2004]. Two complementary approaches to following liquid water in the search for life are therefore to follow the energy and the nutrients [Capone et al., 2006; Hoehler et al., 2007]. Given liquid water is essential for all known life, the definitions of habitable zones within planetary surfaces and interiors are principally based on the ability of the environment to sustain liquid water, and the liquid availability [Szostak, 2005]. The phase of water is determined strictly by the pressure (P) and temperature (T) of its environment. The probability of finding the liquid phase on a planet’s surface will be strongly determined by the occurrence and duration of P-T within the stability regime of liquid, and this will in turn depend on the planet’s distance from the sun, the thickness of its atmosphere and the presence of a greenhouse effect. Figure 1-3 illustrates which solar system planets have the potential for P-T conditions which allow temporary or permanent liquid water on their surface. Within the interior of a rocky planet, P-T generally increases with depth as the overburden and heat flow increase towards the central core (the effects of tidal heating are not considered here as they are not major factors for Earth nor Mars, but would be pertinent for the small moons of giant planets). Hence, the likelihood of subsurface P-T conditions within the liquid regime is determined by the composition and size of the body and its geothermal heat flow. A pressuretemperature diagram can therefore be utilised to identify the range of environments within a planet where conditions are consistent with the stability of
6
Introduction
liquid water. The technique allows the entire phase space of the planet to be presented, so that regions that overlap with the P-T space of liquid water can be identified. This technique will be developed in Chapter 3.
Figure 1-3: Pressures and temperatures within the universe. The phase diagrams of water, carbon dioxide, ammonia and methane are plotted. Triple points for each substance occur at the intersection of the melting, vaporising and sublimating curves for that substance. To highlight the pressure and temperature values of the triple points, asterisks have been placed on the P-T axes. At the triple point, the liquid, gas and solid phases are in equilibrium. ISM = interstellar medium. Environments that fall within the liquid phase space of water, such as the surface of Earth and subsurface Europa ocean, have potential to support life. Research on extremophiles has revealed apparent limits of temperature, water activity and liquid water availability [Jakosky, 2007; Schulze-Makuch and Irwin, 2008a] which are implicated in making some liquid water potentially uninhabitable (for example, MgCl2 bitterns [Bolhuis et al., 2006; Hallsworth et al., 2007; Javor, 1989]). However, no empirical study has comprehensively quantified the fraction of potentially uninhabitable liquid water on Earth or described where
7
these wet inhospitable environments are. Such quantification could be applied to refine the search for life on other planets by focusing the search for liquid water environments on those that are known to be hospitable to organisms on Earth. To this end, a framework will be developed in Chapters 3, 4 and 6 for presenting Earth, Mars, life, and the relevant liquid water chemistries of both planets on a pressure-temperature diagram.
By applying the theoretical and known
temperature, pressure and water activity thresholds of microbial life to the P-T model of Mars, constraints will be placed on the depth and global extent of potentially hospitable martian environments, under the assumption that any potential extra-terrestrial biology would have the same fundamental requirements and tolerances of Earth’s biosphere.
1.2.
Approach 2 – constraining near-surface habitability
One significant limitation of approach 1 to estimating the extent of a potential biosphere on Mars, is that there will be limited ability to probe the deep subsurface of the planet within the next several decades. Probing the shallow subsurface however is a much more attainable goal [Paulsen et al., 2011; WimmerSchweingruber et al., 2011]. Models of the deposition and duration of water in the shallow subsurface of Mars through vapour exchange with the atmosphere are now sophisticated [Mellon and Jakosky, 1995; Mellon et al., 2004b; Sizemore et al., 2009] and are consistent with orbital and surface observations of the perennial ice table depth [Boynton et al., 2002b; Maurice et al., 2011; Mellon et al., 2009a]. Above the stability depth of water ice is a layer of regolith, which is metres thick at low latitude to only millimetres thick at high latitude and non-existent at the poles, where water can be seasonally or diurnally stable. Modelling and observations have revealed that even in the warm equatorial latitudes of Mars, variations in thermal conductivity of materials, solar heating conditions (insolation), or the presence of salts can be sufficient to allow transient liquid water environments [Haberle et al., 2001; Hecht, 2002; Kereszturi et al., 2011; Kossacki and Markiewicz, 2004; Mohlmann, 2011; Rennó et al., 2009; Vincendon et al., 2010a]. Hence, the potential for liquid water in near-surface regions of Mars that were identified as potentially habitable in the martian P-T model described above (1.1), can be examined through constraining the response of surface materials to solar heating.
8
Introduction
The key variables in models of surface and shallow subsurface temperature and water variation are albedo (the fraction of incident solar energy in visual wavelengths that is reflected by the surface) and thermal inertia (the ability of a material to conduct and store heat), as together they parameterise the heat absorbed by surface materials and how that heat is distributed [Mellon and Phillips, 2001]. Thermal inertia and albedo have been globally mapped on Mars [Putzig and Mellon, 2007a] and have been used to determine ‘thermophysical units’ – mixtures of materials on the martian surface that have a similar thermal response and hence a similar shallow subsurface environment [Mellon et al., 2000; Palluconi and Kieffer, 1981; Putzig et al., 2005]. Restrictions in the methodology of the approach used by previous studies in defining thermophysical units, as well the release of updated thermal inertia values [Putzig and Mellon, 2007a], have provided scope for a more sophisticated and detailed mapping of martian surface materials. The methodology of algorithmic classification is applied in Chapter 8 to determine classes in thermal inertia and albedo, and produce a more detailed global mapping of mixtures of dust, sand, duricrust, pebbles, rocks, bedrock, and ice on the martian surface. The validity of the map will be explored extensively through comparison to independent datasets, such as visual and spectral observations of the surface. The thermophysical map constrains the thermal behaviour of the surface and establishes where on Mars the physical conditions for the existence of near-surface liquid water and temperatures within the thresholds of life occur. The results provide necessary input to future models of transient water environments and, ultimately, to assessing the habitability of near-surface environments on Mars.
1.3.
Summary
In combination, the two approaches implemented in this thesis provide complementary assessments of the habitability of Mars, as well as informing on the distribution and limitations of life on Earth. The first approach, which comprises the majority of the thesis, makes a global assessment of the occurrence of liquid water and conditions that are hospitable to life within Earth and Mars, comparing the extent of the potential martian biosphere to Earth’s.
This develops the
understanding of life on Earth, strengthens the astrobiological importance of Mars, and motivates the search for life on Mars by quantifying the depths at which potentially hospitable environments occur.
In reality however, the most
9
significant habitable environments at present are those that occur on or near the martian surface as they have the potential to be sampled and explored in the nearfuture. To contribute to the assessment of habitable environments in the shallow subsurface of Mars that were identified through the biosphere model, the spatial distribution of materials with distinct thermal behaviour (dust, sand, etc.) on the martian surface is mapped. The impact of the surface materials on temperature, water, and habitability of the subsurface is also examined. Finally, predictions are made of regions on Mars that, based on the dominant surface materials, have strong potential for supporting diurnal or seasonal hospitable liquid water environments, and can be directly investigated by the next robotic missions to Mars.
1.4.
Overview & structure
This section describes the contents of each chapter and how the two approaches discussed above are interwoven within the structure of this thesis.
1.4.1. Chapter 2- Literature Review & Background: Quantifying Habitability & the Extreme Limits of Life This literature review sets the context for the new methodology developed in this work for quantifying habitability. It begins by reflecting on broad measures of habitability – from the stability of liquid water on a planet’s surface inferred from orbital parameters, to the chemical and energy sustenance required by Earth life. Work is then presented that is directly relevant to the approach of utilising the phase diagram of liquid water – namely, how the temperature and water vapour pressure within an environment determine if water can be present in the liquid phase. Although all known life requires liquid water, there are some liquid water environments that appear to be inhospitable to microbial life. No known examples of active life can survive toxic levels of high and low temperature, low pressure and low water activity. The thresholds of these parameters that may restrict life from liquid water environments on the Earth and Mars are discussed.
10
Introduction
1.4.2. Chapter 3- Phase Diagram of Earth & Water In this chapter the overlap between the phase space of Earth and the phase space of liquid water is determined.
The Earth’s phase space is constructed by
representing all environments within the Earth’s atmosphere, oceans, surface, crust, mantle and core on a pressure-temperature diagram. All liquid water on Earth is similarly represented, through modelling the phase space of briny ocean and permafrost water. The main application of these models is in Chapter 4, however in Chapter 3 they are used to hypothesise the deepest possible liquid water in the Earth based on geophysical models of the deep interior.
From
thermophysical constraints only, the deepest possible liquid water occurs at the high pressure intersection between Earth phase space and the critical point of liquid water. The likelihood of liquid water environments at such depths within the Earth is discussed, as this represents the maximum possible extent of the biosphere.
1.4.3. Chapter 4- Phase Diagram of Earth’s Biosphere Here a phase model is developed for the Earth’s biosphere. The phase space of Earth life is quantified by the pressures and temperatures of all environments known to support active life. The conditions under which life becomes stressed and dormant are also discussed, as these represent environments where life can survive but cannot reproduce. Using the phase model of the biosphere, the fraction of Earth and liquid water environments that are known to be inhabited by life is investigated, to identify where there are P-T environments within both the Earth and the liquid water regime but external to the biosphere.
This reflects an
estimate of liquid water that is potentially inhospitable to life, and hence such conditions may also be considered inhospitable environments on Mars.
To
conclude the chapter the thresholds of life that were presented in Chapter 2 are applied, and used to speculate on where the search for extremophiles should be concentrated to extend the known limits of life on Earth.
11
1.4.4. Chapter 5- Literature Review & Background: Liquid Water on Mars This literature review considers why Mars has the potential for habitable environments, through considering where there is morphological evidence of water, and more generally where the presence of water ice intersects modelled conditions for melting. It opens with discussion of how the low partial pressure of water vapour in the present-day martian atmosphere severely restricts the duration of liquid water on the surface, and hence why any significant reservoirs of liquid water occur subsurface. The plausible states of putative martian liquid water- from thin films and droplets of brine, to seasonal meltwater from frost and more substantial reservoirs of subsurface brine within the crust- are presented. The review sets the stage for why there is a continuing need for modelling the martian subsurface environment to search for liquid water and potentially hospitable environments for life.
1.4.5. Chapter 6- Phase Diagram of Mars, Water & the Potential Martian Biosphere This chapter begins with the construction of a phase model of Mars and the potential martian biosphere. The latter is defined to be the range of environments with pressure and temperature conditions that concurrently: occur on Mars, overlap the liquid regime of martian water, and are within the thresholds of Earth life. The likely salinity of putative liquid water on Mars was discussed in Chapter 5, but its phase space is modelled here. Furthermore, not only are environmental conditions that support active life on Earth considered, but also conditions that are plausibly not inhospitable to life- including high temperature and high pressure environments that have not been explored for life, but are theoretically habitable if liquid water is present. By comparing Mars to liquid water the maximum and likely extents of the martian subsurface hydrosphere is hypothesised. By then comparing the former to the phase space of Earth life, a new formulation of where and at what depth there are environments on Mars potentially hospitable to life is presented.
12
Introduction
1.4.6. Chapter 7- Bridging Chapter: The Influence of Surface Materials on the Subsurface Environment This chapter bridges the two approaches employed in this thesis, transitioning from the global representation of the potential biosphere of Mars, to locating where hospitable conditions for life occur within metres of the martian surface. The second approach of this work is focused on the shallow subsurface as it the only region of the planet that can be directly explored, sampled and tested for liquid water and biology within the foreseeable future. The model in Chapter 6 demonstrated that the biosphere could be encountered at the surface of Mars in warm equatorial regions. In this chapter that result is quantified. The effect of thermal inertia, albedo and the solar insolation environment of the surface on the shallow P-T conditions, the occurrence of water within the soil, and ultimately whether the conditions are hospitable to life is discussed. By developing a simple model for estimating the temperature variations at tens of metres depth, the general influence of different martian surface materials on shallow subsurface temperatures is demonstrated.
1.4.7. Chapter 8- Remote Sensing of Thermal Materials on the Martian Surface This chapter presents the second approach of this thesis – a new methodology for estimating the dominant materials on the surface of Mars, with an aim to constraining
the
shallow
subsurface
pressure,
temperature
and
water
environment. As discussed in Chapter 7, the main parameters influencing this environment are the thermal properties of the surface materials – how they conduct heat and how effectively they isolate the subsurface from the atmosphere. These thermal properties are straightforward to derive from the behaviour of Earth materials and laboratory data of analogue materials under martian conditions, however the prerequisite is a knowledge of the spatial distribution of materials across the martian surface. Accordingly, the techniques of algorithmic classification that have been underutilised in the field of planetary science are employed, to uncover two dimensional classes of pixels within the global maps of martian thermal inertia and albedo.
These classes are then interpreted as
mixtures of dust, sand, duricrust, rocks and ice in varying proportions, and the occurrence of the classes on the surface is mapped and analysed. To test the
13
validity of the map in determining real trends in martian surface materials, the correlation between the map and other independent martian datasets (such as the groundtruthing provided by the rovers) is investigated.
1.4.8. Chapter 9- Synthesis & Conclusions Here the main results from the thesis are presented, and their implications for where the search for life on Mars should be focused are discussed. Surface and shallow subsurface regions within the potential martian biosphere are identified. Future directions for extending the thesis work are highlighted.
1.5.
Publication summary
Some of the material in this thesis has already made a contribution to the published scientific literature. The peer-reviewed papers are as follows: 1. Jones, E. & Lineweaver, C. (2010) To what extent does terrestrial life ‘follow the water’? Astrobiology, 10, 349-361. 2. Jones, E. & Lineweaver, C. (2010). Pressure-temperature phase diagram of the Earth. Astronomical Society of the Pacific Conference Proceedings: Pathways Towards Habitable Planets, 430, 145-151. 3. Jones, E. & Lineweaver, C., Clarke, J. (2011) An extensive phase space for the potential martian biosphere. Astrobiology, 11, 1017-1033. 4. Jones, E. & Lineweaver, C. (2012) Using the phase diagram of liquid water to search for life. Australian Journal of Earth Sciences: Planetary Themed Special Issue, 59 (2), In Press. 5. Jones, E., Mills, F., Doran, B., Caprarelli, G., Clarke, J. (2011) Initial results from a GIS-based study of Mars’ surface. Australian Space Science Conference Series: 10th Conference Proceedings NSSA, 10, 145-160. ISBN 13: 978-0-9775740-4-9. An additional manuscript is in preparation: Jones, E., Caprarelli, G., Mills, F., Doran, B., Clarke, J. Unsupervised classification of the martian surface into TES thermal inertia – albedo units.
14
1.6.
Introduction
Reference material
Some maps which are referred to frequently within the thesis are included below. All martian regions that are discussed in the thesis are labelled in grey, with locations of landers labelled in red. Acronyms and common terms and their definitions are included in the Glossary.
On the next pages: Figure 1-4: Map of martian surface topography from the Mars Orbiter Laser Altimeter [Smith et al., 2001]. Elevation units are in metres relative to the constant gravity surface (geoid). Spatial resolution is 128 pixels/degree or ~ 460 m. More details in Chapter 8. Figure 1-5: Map of martian thermal inertia derived from Thermal Emission Spectrometer observations [Putzig and Mellon, 2007a]. Units are ‘tiu’ which is equivalent to J/m2/K/s1/2. Thermal inertia represents the ability of a material to conduct and store heat and depends predominantly on effective grain size, with high thermal inertia regions displaying larger grains or a higher degree of cementation. Spatial resolution is 20 pixels/degree or ~ 3 km. More details in Chapter 7. Figure 1-6: Map of martian albedo from the Thermal Emission Spectrometer [Putzig and Mellon, 2007a]. Albedo is the fraction of reflected solar radiation in visual wavelengths. Higher albedo materials have higher visual reflectance and so appear brighter. Spatial resolution is 20 pixels/degree or ~ 3 km. More details in Chapter 7. Figure 1-7: Map of major martian geologic units, based on surface age and morphology [Skinner et al., 2006] (details of units given in reference). More details in Chapter 8.
Figure 1-4
15
Introduction
Figure 1-5
16
Figure 1-6
17
Introduction
Figure 1-7
18
19
Table 1-1: Martian geologic epochs. Geologic Epoch
Surface age
Dominant surface processes
estimates Pre-Noachian
4.6- 4.1 Ga
Heavy impact bombardment
Noachian
4.1-3.7 Ga
Extensive liquid water erosion; major flooding; long term standing water; formation of Valles Marineris and other valleys, clay minerals, sedimentary layers
Hesperian
3.7-3.0 Ga
Episodic outbursts of water erosion; acidic water-rock interaction; volcanic activity; formation of outflow channels, sulfate minerals, lava plains
Amazonian
3.0 Ga-500 Ma
Predominantly water poor environments ; extensive glacial/periglacial activity; accumulation of water ice deposits in the mid- to high latitudes; formation of iron oxide minerals
Late Amazonian
500 Ma – present
Melting and sublimation of near surface water ice deposits (except at the poles); formation of polar caps, glaciers, gullies and other lowvolume flow features
20
Introduction
Figure 1-8: Recent obliquity values for Mars from numerical modelling of the martian orbit [Morgan et al., 2011] (Fig. 20). LDM = latitude-dependant mantle. Obliquity has a major affect on the martian climate and can be used to understand whether the features observed on the surface formed under similar climatic conditions to that of Mars today.
21
2. LITERATURE REVIEW & BACKGROUND: QUANTIFYING HABITABILITY & THE EXTREME LIMITS OF LIFE
‘The ocean is a desert with its life underground and the perfect disguise above. Under the cities, lies a heart made of ground but the humans will give no love.’ - America, ‘Horse With No Name’
22
Literature Review & Background: Quantifying Habitability & the Extreme Limits of Life
This chapter provides the context for the primary research question of this thesis: are there habitable environments on present-day Mars?
Liquid water is a
fundamental requirement for all known life on Earth and the work in this thesis focuses on advancing the search for liquid water on Mars. Therefore, the most relevant attributes of the biosphere for this work are the limitations on life within liquid water environments. In other words, when liquid water is present, what other conditions are also necessary for that liquid to be habitable. Previous work on the habitability of Earth’s liquid water environments will be discussed, with the review focusing on three fundamental constraints: temperature, pressure, and the activity of the liquid. These constraints are employed in later chapters to estimate the extent of habitable environments on the Earth and on Mars. In addition to liquid water, all life on Earth requires organic compounds and a source of energy. Although these parameters are largely beyond the scope of this thesis, life’s requirements for energy and nutrients will also be briefly discussed, as will some abstractions to the forms of life which could plausibly occur on Mars, and the resources that such life would likely utilize.
2.1.
First order: habitable zone around a star
For centuries, the position of a planet in space has been recognised as influencing its habitability [Gonzalez, 2005]. For the last four decades the occurrence of the most significant requirement for biology within solar systems- liquid water- has been discussed through the concept of circumstellar habitable zones (CHZ). The CHZ is defined as the range of distances from a central star where liquid water would be stable on a planetary surface [Hart, 1979]. The location and width of such a region is dependent not only on stellar properties (such as luminosity and stellar activity, Figure 2-1) but is also a complex function of the planetary mass and the presence and composition of its atmosphere [Badescu, 2011a] (for example, the presence of a CO2 greenhouse). Around Sol, the habitable zone for an Earthlike planet extends from 0.95-1.37 AU [Kasting et al., 1993]. The inner edge of the habitable zone is determined by the complete loss of oceans through photolysis of water and hydrodynamic escape [Doyle et al., 1998]. The outer edge is determined by CO2 condensation, which increases the surface albedo and hence lowers its temperature [Doyle et al., 1998; Whitmire et al., 1991]. Other factors can also make an environment habitable by providing the necessary energy to maintain water in
23
the liquid phase, resulting in potentially habitable environments outside of the traditional CHZ. For example several solar system moons, such as Europa at 5.23 AU and Enceladus at 9.54 AU, may be sufficiently tidally heated by their host planets (Jupiter and Saturn, respectively) that they are able to maintain subsurface liquid oceans [Reynolds et al., 1983, 1987; Roberts and Nimmo, 2008; Squyres et al., 1983].
Figure 2-1: Extrasolar planets around solar-mass stars [Schneider et al., 2011]. To date 498 planets have been found orbiting stars with a similar mass to our sun, many of which orbit within the region that is habitable in our solar system (blue line). Whether these planets lie within their own habitable region however, depends strongly on the host star, the planet’s mass, eccentricity and atmosphere and on the definition of habitability, particularly whether long term liquid water stability on the surface is required. Within our solar system, Mars is beyond the outer boundary of the habitable zone (orbiting at ~1.5 AU), as evidenced by the instability of liquid water on the majority of its surface and the presence of extensive CO2 condensation, although it was likely within the habitable zone > 500 Ma ago due to higher average obliquity [Franck et al., 2000]. Wet planets, such as the Earth, may be able to evolve to Marslike conditions whilst still maintaining habitable environments, even as surface
24
Literature Review & Background: Quantifying Habitability & the Extreme Limits of Life
water becomes limited. Such ‘land planets’ can retain a hydrological cycle with the transport of water between the atmosphere and the surface through a much more extensive orbital region than is suitable for ‘ocean planets’ [Abe et al., 2011; Kodama et al., 2011]. Given the developments in understanding martian surface water, its relationship with the atmosphere, its stability on diurnal or seasonal timescales [Haberle et al., 2001; Smith et al., 2009b], and the abundance of subzero liquid water on the Earth [Price, 2000], the concept of a habitable zone around a star should be expanded to include worlds with limited surface water or only temporary forays into the traditional habitable zone of liquid water stability. Not only would this be relevant for Mars, but it would allow us to reconsider whether other extrasolar planets (Figure 2-2) could lie within their respective solar habitable zones [Kasting, 2011; Selsis et al., 2007] (as has been suggested for planets in the HD 85512 and Kepler-16 systems [Kaltenegger et al., 2012; Quarles et al., 2012]). The various forms of liquid water in the solar system, such as brines and thin films, will be discussed in detail in section 2.5.
Figure 2-2: Exoplanets which may enter their habitable zones, defined by the potential for liquid water on their surface [Kane and Gelino, 2011, 2012]. Orbital parameters such as period and eccentricity (circularisation of the orbit) affect potential habitability. The size of the points linearly increases with the percentage of time spent within the habitable zone, defined by the range of circumstellar distances
25
from a star within which a planet could have liquid water on its surface, given a dense enough atmosphere. Image credit: Habitable Zone Gallery1.
2.2.
Fluids for life
All known life on Earth uses liquid water as a solvent – requiring it during some phase of its life cycle [Rothschild and Mancinelli, 2001]. Consequently liquid water, as a necessary (but not sufficient) condition for life, has long provided the first criterion in establishing the boundaries of habitability. Liquid water was a logical choice for life on Earth [Schulze-Makuch and Irwin, 2008b], as it is the most abundant fluid on Earth [Baross et al., 2007] and is plausibly one of the most common fluids in the universe given the abundances of H, O and H2O ice [van Dishoeck, 2011; Herbst, 1995; Mottl et al., 2007]. Furthermore, the temperature and pressure boundary conditions for liquid water are well matched by the surface and subsurface conditions on Earth. Hence, liquid water is stable in a wide variety of Earth environments (Chapter 3). Many features of liquid water are utilized by Earth’s organic chemistry, such as the solubility of organic compounds (enhanced by the polar nature of water) and hydrogen bonding. However, whether liquid water is necessary for life, or whether a liquid medium is even required, is unknown and can only be speculated [Kasting, 2010]. Liquid water is not necessarily the most abundant fluid in other planetary systems [Baross et al., 2007] and the presence of alternate chemistries on other planets may necessitate a different choice of solvent for biology [Bains, 2004].
For
example, the cold conditions on the martian surface are better suited to the liquid regime of carbon dioxide than that of water [Musselwhite et al., 2001]. Titan’s surface supports lakes, most likely of methane or ethane [Stofan et al., 2007], with a possible ammonium ocean in its interior [Fortes, 2000; Grindrod et al., 2008]. Ammonia itself has been suggested as a candidate solvent for alternative biochemistries as it is abundant and like water it dissolves many organic compounds, but unlike water it remains liquid over a wider range of temperatures, particularly at higher pressures [Benner et al., 2004]. Sulfuric acid is another candidate polar solvent for life [Benner et al., 2004], is present both as a vapour and a liquid in the venusian atmosphere [Kolodner and Steffes, 1998; McGouldrick et al., 2010], and 1
http:\\www.hzgallery.org
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Literature Review & Background: Quantifying Habitability & the Extreme Limits of Life
supports the growth of extreme acidophiles in pH ~ 1.3 sulfuric acid [Vasquez and Espejo, 1997]. Considering however that the absence of water and extremes of temperature beyond a certain threshold are known to prevent growth and survival of all organisms on Earth, it is difficult to envision a biological system that can exist without liquid water, or in temperature conditions that exclude the biochemistry of Earth (such as within the liquid ammonia regime). The absence of a robust definition for life [Cleland and Chyba, 2002] further increases the challenge in envisioning a biology totally dissimilar to our own. Hence, this work will continue under the assumption that the best current constraints on habitability are the known thresholds and requirements of life, the most fundamental of which is liquid water [Chyba and Hand, 2005].
2.3.
Energy & nutrients
All life on Earth requires chemical energy to fuel metabolism [Gimenez, 2010; Morowitz, 1968] and drive metabolic processes such as the synthesis of adenosine triphosphate (ATP) [Conrad and Nealson, 2001; Hoehler, 2004].
Many surface
organisms harness photons to provide energy through photosynthesis, while others use the organic molecules produced by these organisms (such as carbohydrates) as their energy source [Schulze-Makuch and Irwin, 2008a]. Subsurface life utilises chemical, rather than solar energy, for example from the oxidation of volcanic gases like hydrogen (H2), hydrogen sulphide (H2S), and methane (CH4) [Gold, 2001]. It is difficult to extrapolate the energy requirements of Earth life to that of other life forms, however it seems fundamental that in order to be habitable the available energy supply in an environment must be comparable to the biological energy requirements of known organisms [Hoehler, 2004]. In other words, an environment becomes uninhabitable when the energetic costs of procuring water and nutrients exceed the extractable energy [Shock and Holland, 2007].
An absence of energy will result in the cessation of metabolism and
biological function [Hoehler, 2004], and limited energy will restrict the complexity of the biological systems that can develop [Hoehler et al., 2007]. Without a broad range of models for biochemistry, it is quite plausible that potentially habitable environments on other planets will be overlooked [Nealson, 1997; Nealson et al., 2002] as the primary sources of energy on other worlds may be dissimilar to those
27
available on Earth, and consequently it is possible that very different biospheres may have evolved to utilize them [Schulze-Makuch and Irwin, 2008a]. Microbial systems on Earth that exist in the subsurface and are not reliant on surface photosynthetic life [Deming and Baross, 1993] provide the strongest analogy to candidate martian life. These organisms are chemosynthetic, utilizing inorganic compounds as an energy source, rather than the sugars and carbohydrates used by other organisms. A number of candidate energy sources have been suggested as being capable of supporting this form of subsurface martian biosphere.
These include geochemical energy released through
exothermic water-rock interactions and water-volcanic gas reactions (analogous to the energy used by life in hot springs, deep ocean vents and deep crustal environments [McCollom and Shock, 1997; Pedersen, 2010]), through hydrothermal systems that may have been present (or may still be present) on Mars [Golden et al., 2010]. Energy sources are still available in the low temperature conditions that dominate the martian surface today, through low volumes of liquid water interacting with basaltic rock [Link et al., 2005]. Life would obtain energy through exothermic redox reactions [Boston et al., 1992], with aerobic life (requiring oxygen) employing oxidation reactions and utilizing CO2 or O2 as a source of oxygen, and anaerobic life performing reduction reactions. Chemical species such as Fe(II), Fe(III), and Mn(III) can be oxidized to release energy [Kelley et al., 2002], analogously to oxidizing bacteria in the deep basaltic aquifers on Earth [Stevens and McKinley, 1995]. Also SO4(II) can be reduced at low temperatures [Zolotov, 2004] and is utilised by bacteria in terrestrial permafrost [Gilichinsky et al., 2007] as are nitrates [Nealson, 1997]. Minerals such as olivine and pyroxene can also be reactants in oxidation reactions [Link et al., 2005; Varnes et al., 2003]. Alternatively, in regions of the surface and shallow subsurface on Mars that have access to thin films of liquid water or brines (such as at the Phoenix Landing site [Rennó et al., 2009; Searls et al., 2010]), the energy to drive metabolism could be provided by sunlight and redox pairs of perchlorate and iron [Stoker et al., 2010]. All of the aforementioned chemical species have been detected on Mars (Chapter 5), with the exception of nitrates [Jakosky, 2007]. Any life on Mars would also need access to elements which can provide the building blocks of biological structures,
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Literature Review & Background: Quantifying Habitability & the Extreme Limits of Life
such as nucleic acids and proteins [Jakosky and Shock, 1998]. On Earth, C, H, N, O, P and S are used to build biomass, with carbon and nitrogen being the most critical for biological molecules [Capone et al., 2006; Schuur, 2011] (although some organisms may be able to exchange phosphorus for arsenic [Wolfe-Simon et al., 2011]). Sources of all these nutrients are also available on Mars [Nealson, 1997]. In particular, inorganic carbon could be obtained from CO2 [Kelley et al., 2002] or potentially CH4 (if methane is present [Zahnle et al., 2011]) to build biomass. Hence, in the warmer, wetter, and more geologically active martian past, it seems likely that sufficient chemical energy and nutrients were available to support a significant biosphere.
Furthermore, it is plausible that a smaller volume of
subsurface biota may be sustained today [Jakosky and Shock, 1998; Varnes et al., 2003].
2.4.
Phase requirements of liquid water
The phase diagram of water describes the range of pressures (P) and temperatures (T) at which water molecules exist as a solid, liquid or gas [Silberberg, 2003]. Three curves divide these phases: the vapourisation curve (transition between gas and liquid), the freezing curve (transition between solid and liquid) and the sublimation curve (transition between solid and gas). These curves represent phase transitions and describe the P-T conditions at which the two phases are in equilibrium. If the P-T values lie within a region, the equilibrium has shifted to increase the presence of whatever phase is described by that region. At the triple point, all three phases are in equilibrium. The boundary curves between the three phases of water are given by the Clausius-Clapeyron equation [Lide and Haynes, 2009] (see reference for numerical values).
The critical point describes the
termination of the vapourisation curve and is the temperature at which the densities of the liquid and vapour phase of a substance become equal. Supercritical water is not considered as it has dramatically different physical properties to sub-critical liquid water (including density and viscosity [Benner et al., 2004]), and there is no evidence that such a solvent could be hospitable to life [Baross et al., 2007; Schulze-Makuch and Irwin, 2006]. In addition to these three primary phases there are many other phases of water ice (the current count is 13 [Vega et al., 2005]) that occur at extreme densities [Bridgman, 1937] and low
29
temperature [Poole et al., 1993], and also ultradense phases at high pressures [French et al., 2009]. They are beyond the complexity required for this work.
2.4.1. Interpreting the water phase diagram The phase diagram of water is shown in Figure 2-3. The pressure values are partial pressures, Ppart, the pressure exerted by water molecules in the gaseous phase. Some fraction of molecules are in the gaseous phase in all regions of the phase diagram, unless the partial pressure is zero due to the species being absent. In an atmosphere the water vapour pressure will generally be some fraction of the total pressure Ptot as there will be other species also in the gaseous phase. The maximum vapour pressure of water in an atmosphere (lying on the vapour curve) is the saturation pressure Psat. This pressure is temperature dependent, with warmer atmosphere able to hold more water vapour. Boiling occurs when Ppart equals the surrounding atmospheric pressure Ptot at any temperature on the vapour curve. For example, at sea level on Earth Ptot is 105 Pa (point C in Figure 2-3) and hence boiling occurs at 100 ˚C. The phase diagram indicates the stability of water in different P-T conditions. For example, the Ppart of water in typical indoor conditions is 3 × 103 Pa at 25 ˚C, shown by point A in Figure 2-3. Its location on the phase diagram shows that water is in disequilibrium under these conditions and will evaporate to increase the presence of the vapour phase. Thus a glass of water left uncovered indoors at T > 0 ˚C and Ppart within the vapour region, will slowly evaporate until the liquid is gone and the atmosphere is saturated (water stability is examined in more detail in Chapter 5). Contrastingly in a cold and dry environment such as the Antarctic Dry Valleys, at point B, the Ppart of water vapour is less than 20 Pa at -30 ˚C (Terra Nova Bay [Tomasi et al., 2004]). This is very close to the equilibrium point of vapour and ice (30 Pa at -30 ˚C) and is outside of the stability region of pure liquid water.
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Literature Review & Background: Quantifying Habitability & the Extreme Limits of Life
Figure 2-3: Partial pressures of water on Earth. The green crosshatching indicates water vapour conditions in Earth’s atmosphere. The phase space of pure liquid water is shaded blue. A = a typical indoors environment; B = dry Antarctic conditions; C = boiling condition at sea-level; and D = a typical partial pressure at zero elevation on Mars. Only extremely dry environments will fall outside the liquid regime and have Ppart less than the triple point of water (indicating the liquid phase is not stable). For comparison, the green and red circles indicate the Ptot at sea-level/zero elevation on Earth/Mars respectively.
2.4.2. Freezing point depression and boiling point elevation The addition of a solute to water will generally increase the phase space of liquid water, allowing water to remain in the liquid state for a larger range of temperatures and pressures. Depression of the freezing point and elevation of the boiling point occur due to the presence of the solute decreasing the vapour pressure Ppart over the solution, relative to the Ppart over the pure substance at the same temperature. This shifts the Clausius-Clapeyron curves so that equilibrium between phase changes is achieved at lower vapour pressures.
Hence, the
vapourisation curve for the solution will lie at a lower pressure than that of the pure solvent at any given temperature.
The melting curve will lie at lower
temperatures for any given pressure. Freezing point depression and boiling point
31
elevation are collegiative properties, meaning they are dependent only on the number of solute particles (moles) rather than the chemical identity of the constituents. Raoult’s law describes the ideal value of temperature elevation and depression: Equation 2-1 ∆Tf =-i×kf×M Equation 2-2 ∆Tb = i×kb×M Where M = solution molarity (concentration of solute particles); kf = molal freezing point depression constant (1.86 ˚C/M for water); kb = molal boiling point elevation constant (0.512 ˚C/M for water); i = van’t Hoff factor [Silberberg, 2003]. The van’t Hoff factor accounts for the behaviour of a non-ideal solution, where the solute is highly volatile or strongly electrolytic (such as a saline solution). In the latter case, ionic interactions within the solution reduce the effective concentration of the ions, increasing the vapour pressure over the solution, and reducing the magnitude of the expected temperature variation from Equation 2-1 and Equation 2-2. For an ideal solution the van’t Hoff factor is equal to the dissociation number of the solute (the number of ions in solution). For a non-ideal solution, the van’t Hoff factor is determined experimentally as the ratio of the colligative property in the electrolyte solute to the expected value for a non-electrolyte solution. van’t Hoff factors approach their ideal values as the electrolyte solution becomes more dilute and the distance between ions increases [Silberberg, 2003]. Table 2-1, Table 2-2 and Table 2-3 provide the magnitude of the temperature shift required to alter the phase diagram of water to that of an aqueous solution of water with solute. These values are only estimates however as ideal van’t Hoff factors were used. Table 2-1: Freezing point depression (˚C) for an ideal solution. Dissociation number 1 2 3 4
0.5 -0.93 -1.86 -2.79 -3.72
Molarity of solvent (mol/kg) 1 2 5 -1.86 -3.72 -9.3 -3.72 -7.44 -18.6 -5.58 -11.16 -27.9 -7.44 -14.88 -37.2
10 -18.6 -37.2 -55.8 -74.4
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Literature Review & Background: Quantifying Habitability & the Extreme Limits of Life
Table 2-2: Boiling point elevation (˚C) for an ideal solution. Dissociation number 1 2 3 4
0.5 +0.255 +0.510 +0.765 +1.02
Molarity of solvent (mol/kg) 1 2 5 + 0.512 + 1.024 + 2.56 +1.024 +2.048 +5.120 +1.536 +3.072 +7.680 +2.048 +4.096 +10.240
10 + 5.12 +10.240 +15.360 +20.480
Table 2-3: Properties of common salts. Common salts NaCl CaCl2 MgCl2 KCl MgSO4 Na2SO4 CaSO4 K2SO4 FeSO4 NaNO2 KNO3 Mg(NO3)2 Ca(NO3)2 Na2CO3 CaCO3 K2CO3 FeCO3 MgCO3 Mg(ClO4)2 NaClO4
Molar mass Dissociation (g/mol) number 58.4430 2 110.9848 3 95.2115 3 74.5515 2 120.3683 2 142.0428 3 136.1416 2 174.2598 3 151.9084 2 84.9948 3 101.1033 2 148.3151 3 164.0884 3 105.9886 2 100.0875 2 138.2057 2 115.8543 2 84.3141 2 223.2069 3 122.4407 2
The effect of substances which expand the phase space of liquid water on the habitability of that liquid water for life will be discussed in section 2.5 below.
33
2.5.
Water activity
Life is found in many environments where it tolerates low water content and survives by adapting to water desiccation beyond the tolerance of most organisms. Halophiles (‘salt loving’ microorganisms), osmophiles (‘osmotic pressure loving’) and frequently psychrophiles (‘cold loving’) all populate environments where the water has a minimal availability for biological processes – quantified by the water activity (examples in Table 2-4; [Grant, 2004; Ponder et al., 2008]). Water activity describes the equilibrium amount of water available for reactions. It is defined as the ratio of the water vapour pressure over the solution to the vapour pressure over pure water at the same temperature [Blandamer et al., 2005].
Equation 2-3:
ܽ௪ ൌ
ೌೝ̴ೞೠ ೌೝ̴ೠೝ
The activity of pure water is 1. An activity of 0 is the theoretical minimum, indicating a total absence of free water molecules. Low water activity can be caused by interactions with solutes, surfaces, or lowered temperatures (Figure 2-4). The water activity of solutions below their freezing point is dependent only on temperature, rather than solute nature or concentration [Koop, 2002], and hence all aqueous solutions with the same freezing point will therefore have the same water activity below that temperature [Koop, 2004]. Table 2-4: Low water activity environments supporting active life Environment High sugar foods (e.g.
Min. aw of
Temp. of
environment
environment (˚C)
0.62
20-40
chocolate syrup) Great Salt Lake, Utah
Reference [Grant, 2004; Restaino et al., 1983]
0.88
0-40
(35% NaCl)
[Brandt et al., 2001; Butinar et al., 2005; Wainø et al., 2000; Zeikus et al., 1983]
Dead Sea, Israel (16%
0.69
15-40
MgCl2) Brine lenses in
1983] 0.9
-15 to 20
permafrost (10% NaCl) Asphalt lake, Trinidad
[Bolhuis et al., 2006; Oren, [Gilichinsky et al., 2003; Shcherbakova et al., 2004]
0.49*
30-60
[Schulze-Makuch et al.,
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Literature Review & Background: Quantifying Habitability & the Extreme Limits of Life
and Tabago
2011]
Ca-sulphate crusts,
0.3 (0.6 for
Atacama desert (> 70%
growth)
20-40
[Wierzchos et al., 2011]
-30 to 0
[Meyer et al., 1962; Siegel et
gypsum CaSO4) Don Juan Pond,
0.45*
Antarctic (45% CaCl2)
al., 1979]
*Requires confirmation.
Low water activity conditions occur in aqueous environments with high solute concentrations (such as saline or sugar saturation [Chirife and Buera, 1996]), subzero temperatures or extremely dry conditions (for example, within permafrost [Ayala-del-Río et al., 2010]). Desiccated environments can also occur where the substrate is impermeable to surface water (such as in tar pits [Kim and Crowley, 2007]). The behaviour of water activity with the addition of a solute and in low temperature conditions is described in sections 2.5.1 and 2.5.2 respectively.
Figure 2-4: Schematic of the increase in water activity with the addition of water to a dry material and the thickness of the liquid film [Seiler, 1976]. Actual values must be determined experimentally for a given material.
35
2.5.1. Water activity of solutions (solute effect) The presence of solutes decreases the activity of liquid water below unity, as the interaction of water molecules with the solute reduces the fraction of molecules with enough energy to enter the vapour phase [Silberberg, 2003]. This discussion will be focused on the water activity of brines (rather than other solutes such as sugar) as salts are the most common solute encountered in Earth’s waters [Bowen, 1986] and are expected to occur in putative martian water [Jagoutz, 2006]. The environmental stressors of low temperature and low water activity often coincide. Subzero water remains liquid due to high salinity content resulting in freezing point depression and lower water activity.
Furthermore, salinity
increases and water activity decreases during freezing, as the salts are excluded from the solid state and hence increase in concentration within the liquid [Vrbka and Jungwirth, 2005].
The water activity of saturated salt solutions can be
modelled by:
Equation 2-4
݈݊ሺܽ௪ ሻ ൌ
భ ்
െ ݇ଶ
where k1 and k2 are constants that depend on the salts present and are obtained experimentally [Kitic et al., 1986]. Equation 2-4 assumes that the moisture content of the system remains constant [Sahin and Sumnu, 2006]. At temperatures above the freezing point but below 0 ˚C, the water activity of saline solutions generally decreases steeply with decreasing temperature, however the analytic relationship varies for each salt [Morillon and Debeaufort, 1999].
Below their eutectic
temperatures, the water activity can be calculated analytically. For reference, the activity of supercooled liquid water in equilibrium with ice reaches 0.6 at its minimum temperature of -48 ˚C (Table 5 of [Fennema, 1996]). The water activity and eutectic temperatures of some brines are shown in Table 2-5. Some of the salts in Table 2-5 are present in Earth seawater; the remainder are found in inland salt lakes and groundwater [Benison and LaClair, 2003; Krumgalz and Millero, 1982; Nissenbaum, 1975].
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Literature Review & Background: Quantifying Habitability & the Extreme Limits of Life
2.5.2. Thin films (temperature effect) Thin films of unfrozen water tens of angstroms thick can exist as coatings on ice or grains [Price, 2000] at temperatures as low -100 ˚C [Pearson and Derbyshire, 1974] or potentially as low as -130 ˚C [Mazur, 1980]. On Earth, thin films of liquid are ubiquitous with arctic environments [Anderson and Tice, 1972].
Their low
temperature liquid behaviour occurs primarily due to the physical bonding between water and the adjacent soil or ice surface through van der Waals forces [Mohlmann, 2008] causing a significant change in the energy state of the water molecules [Davis, 2001]. This serves to expand the range of temperatures at which that water can remain unfrozen [Cary and Mayland, 1972; Yong et al., 1979]. The amount of unfrozen water in soil depends strongly on the soil content and vapour pressure of water but also on the fraction of ice, with the amount of liquid film increasing with increasing ice content [Nakano et al., 1982]. The mass of unfrozen water in Earth permafrost varies with soil type, but measurements fall within ~ 15 wt% at -10 ˚C , ~ 3.5-15 wt% at -20 ˚C, 2.5-15 wt% at -25 ˚C, and ~ 2 wt% at -30 ˚C [Cary and Mayland, 1972; Dillon and Andersland, 1966; Tice et al., 1981; Yong et al., 1979; Yoshikawa and Overduin, 2005] (Figure 2-5). The amount of frozen water is not static however, as the liquid undergoes mass flow and thermal gradients which cause continual freezing until the vapour pressure is sufficiently reduced [Nakano et al., 1984]. Experimental data summarised in [Mohlmann, 2005] showed that the activity of a thin film drops below 0.6 at -20 ˚C [Jakosky et al., 2003]. The decrease of activity of the film with further decreasing temperatures is not well known.
37
Table 2-5: Eutectic temperatures and water activity of saturated aqueous solutions of some Earth salts. Salt* NaCl CaCl2
Eutectic temp. ˚C -21 -54
MgCl2
-33
KCl Na2SO4 CaSO4 MgSO4 K2SO4 FeSO4
-11 -5 -0.06 -5 -1.6 -68
NaNO3 -18 Mg(NO3)2 -32 KNO2
-40.2
CaCO3 MgCO3 K2CO3 NaBr
-1.9 -0.3
CaBr2 MgBr2 KBr *Grey
-29
-42.7 -13
Estimated Reference min. aw 0.76 [Grant, 2004; Greenspan, 1977] 0.26 [Bui et al., 2003; Morillon and Debeaufort, 1999; Siegel et al., 1979; Siegel and Roberts, 1966] 0.34 [Greenspan, 1977; Morillon and Debeaufort, 1999; Siegel and Roberts, 1966] 0.89 [Greenspan, 1977; Steele et al., 1969] [Post, 1977; Williams and Robinson, 1981] [Marion and Farren, 1999] 0.1 [Pillay et al., 2005; Zhang and Chan, 2000] 0.99 [Greenspan, 1977] 0.11 [Chevrier et al., 2009; Marion and Farren, 1999; Xu et al., 2009] 0.79 [Greenspan, 1977; van der Ham et al., 1998] 0.60 [Greenspan, 1977; Morillon and Debeaufort, 1999] 0.9 [Greenspan, 1977; Morillon and Debeaufort, 1999] [Marion, 2001] [Marion, 2001] 0.43 [Greenspan, 1977] 0.64 [Greenspan, 1977; Morillon and Debeaufort, 1999; Negi and Anand, 2004; Tang et al., 2003] 0.22 [Greenspan, 1977] [Morillon and Debeaufort, 1999] 0.86 [Greenspan, 1977; Steele et al., 1969]
shading indicates values not found in the literature.
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Literature Review & Background: Quantifying Habitability & the Extreme Limits of Life
Figure 2-5: Liquid water content of soil (am) determined by water activity (aw) for temperatures of 200 K (upper curve), 220 K, 240 K and 260 K [Mohlmann, 2008] (Fig. 8). Vertical axis is the ratio of the mass of H 2O in the soil compared to dry soil. The volume of liquid water depends on temperature, humidity and the physical interaction between the water and the surface.
2.5.3. Implications for life Low water activity appears to be a fundamental limit for life. There is an upper limit to the desiccation stress (a result of low water activity) which organisms can survive, as extreme levels result in irreversible destruction or structural changes to DNA, lipids and proteins, resulting in death [Baross et al., 2007]. For example, decreasing water activity by 80% decreased the cell water content of yeast to 1% of normal and resulted in death [Mazur, 1980]. Studies into brine environments where life is present only in a dormant state- meaning its metabolism is unmeasurably low, no reproduction is occurring but basic cell repair may be underway [Wharton, 2002]- indicate that the low activity of the liquid water present may have excluded active life. For example, the CaCl2 concentrated brine in Don Juan Pond [Siegel et al., 1979] with activity of 0.48 [Matsubaya et al., 1979] and the MgCl2 saturated bitterns [Bolhuis et al., 2006] with activity of 0.3 do not support active life. Currently an ܽ௪ of 0.6 is recognised as the lower limit for active life [Grant, 2004]. Below this cutoff the availability of free water molecules may not be sufficient to sustain biochemical reactions and hence metabolising life [Southam et al., 2007]. The tolerance of some organisms to low water activity (Table 2-4) enables them to populate extreme environments in the crust and to be
39
potentially adaptable to the conditions on other planets such as Mars (see Chapter 6). Chaotropic substances disrupt the intermolecular forces between water molecules, and destroy complex molecules such as DNA and proteins [Doran et al., 2010] that are essential for survival. There are a number of organisms which have been identified as chemotropically limited with the addition of agents to alleviate chaotropicity allowing them to grow at lower water activity [Williams and Hallsworth, 2009]. This lead to the suggestion that chaotropicity instead provides a more fundamental limit for extreme life [Williams and Hallsworth, 2009]. The magnitude of the chaotropic effect and water activity depression however, both depend strongly on the number of species present in solution [Clark and van Hart, 1981; Williams and Hallsworth, 2009]. It has not yet been demonstrated that the effect of denaturing biomolecules occurs in all environments where life exists at the extreme limit of water activity [Hallsworth et al., 2007].
However, it is
certainly valid that the basis of a low water activity cutoff of 0.6 is currently not well understood [Chin et al., 2010]. Given that low water activity is currently recognised as a limit on Earth’s biosphere, and a chaotropic limit has not been widely verified, this review will continue by focusing on water activity and how it can restrict life from some liquid water environments on Earth and potentially other planets. The activity of subzero liquid water has been found to be an important and possibly limiting factor in whether water in cold environments can be utilised by biology, with the current known threshold for active life at a water activity of 0.6 [Grant, 2004]. Given this limit and the activity of brines composed of common salts (Table 2-5), it is possible that there are liquid water environments on Earth with a freezing point below the current minimum temperature known for life of -20˚C [Junge et al., 2006] but with an activity above 0.6. The most likely low temperature brine, found in the pore spaces of sea ice and in permafrost, is NaCl. NaCl has an activity suitable for life, but with its moderate eutectic of -21 ˚C it cannot provide a significant expansion to the temperature space of life. Other brines such as MgCl2 (found in the dead sea [Mullakhanbhai and Larsen, 1975]), CaCl2 (hypersaline lakes in Antarctica [Siegel et al., 1979]), perchlorates (groundwater contaminant from
40
Literature Review & Background: Quantifying Habitability & the Extreme Limits of Life
fertilizer [Gu and Coates, 2006]), FeSO4 (groundwater contaminant from mining [Rose and Ghazi, 1997]), and magnesium bromide (the dead sea [Al-khlaifat, 2008]), all provide significant freezing point depressions but have extremely low (or unknown) water activities below the threshold for life. Two brines, MgNO3 and NaBr, may be candidates for low temperature life below -20 ˚C, however neither are common in the natural environment. Magnesium nitrate is incorporated in stone masonry and dissolves in rainwater [Charola, 2000], and is found as the mineral nitromagnesite in caves [Whisonant, 2001] and, hence, potentially in cave water systems. Sodium bromide occurs as a groundwater contaminant as it is used in the mining industry [Qian and Jiexue, 1994]. A concentrated solution of these brines could have a freezing point depression of -30 ˚C and an activity suitable for life, however it is unlikely that they would occur in such low temperature conditions (or in significant volumes and concentrations) in the Earth environment. Hence in reality, -21 ˚C in concentrated NaCl brine may be the lowest temperature liquid brine environment for active life on Earth due to the absence of suitable freezing point depressors for lower temperatures, if the limit of 0.6 for life is fundamental. As liquid water freezes eventually only a thin layer several molecules thick of unfrozen water will remain and any microbes present within the original liquid veins will be restricted to the thin film environment. These films of liquid water on the surface of ice or coating soil or dust grains (2.5.2) provide a significant volume of liquid water in terrestrial frozen soils down to temperatures of at least -25 ˚C [Anderson and Tice, 1972]. Thin films are a habitat that can sustain microbial activity in glacial ice [Price, 2000] as microbes have been found actively growing and metabolising within them at temperatures of -17 ˚C [Rivkina et al., 2000]. Diffusion gradients can occur within the fluid through the movement of ions- a feature that is important for biology to maintain its supply of nutrients [Mohlmann, 2009a] (2.3). However, it is unclear whether this process of nutrient delivery through diffusion is limited by the size of the films [Jakosky et al., 2003], or if microbes will simply be restricted from utilising the water when the activity falls below 0.6. If this activity limit is fundamental, then thin films colder than -20 ˚C will be uninhabitable by life. The adaptations which psychrophiles possess to tolerate cold conditions are generally accompanied by desiccation and low water
41
activity tolerance [Gunde-Cimerman et al., 2003; Thomas and Dieckmann, 2002]. The reverse is not true however, as from Table 2-4 the majority of organisms in low water activity environments do not experience extremes in temperature. Extremely cold tolerant organisms [Panikov et al., 2006] may therefore be excellent candidates to test the limits to which microbes can utilise subzero brine and thin film environments.
2.6.
Pressure
2.6.1. High threshold Life requires liquid water and hence is restricted to the range of pressure and temperatures that allow water to remain a liquid. Within this region of phase space however, are there other constraints that restrict life from high pressures? Barophiles (‘pressure loving’) have been found as deep as 6 km in the crust [Szewzyk et al., 1994], to the base of the oceans at 11 km [Kato et al., 1998], and 2 km within subocean sediments [Roussel et al., 2008]. These organisms are actively growing in pressure conditions ~ 103 times that found at the surface. Barophiles are generally hyperthermophiles (‘extreme heat loving’) with the average temperature at ~ 4 km depth on Earth being 120 ˚C (deep sea barophiles not from hydrothermal vents are an exception to this trend [Yayanos et al., 1981]). There is most likely a link between the physiological adaptations to both extremes [Holden and Baross, 1995; Robb and Clark, 1999].
Hence, if there is a fundamental
temperature limit to life (discussed below) it may follow that there is an upper limit to the pressures life can endure.
Experimental studies have exposed
barophiles to pressures encountered as deep as 50 km within the crust, finding that the organisms remain alive and viable [Margosch et al., 2006; Sharma et al., 2002]. Additionally, a number of organisms have been found to be able to tolerate high temperatures (> 100 ˚C) only under high pressure conditions, with the higher pressure required to achieve growth under the extreme temperature [Deming and Baross, 1993; Pledger et al., 1994].
High pressures also resulted in a longer
tolerance to lethal temperatures [Marteinsson et al., 1997]. A number of enzymes and proteins have been identified as deactivating under high pressure, but others are stabilised [Michels et al., 1996]. Together, these studies indicate that currently there is no known upper limit to the pressures that life can survive. The more
42
Literature Review & Background: Quantifying Habitability & the Extreme Limits of Life
fundamental limiting factors for life which will determine the base of the biosphere are water availability, high temperature, energy and nutrients.
2.6.2. Low threshold Life has been found throughout Earth’s atmosphere to an altitude of 77 km above the surface [Imshenetsky et al., 1978]. Organisms from these environments can be classified into two types – spores or spore forming bacteria (such as bacillus simplex [Smith et al., 2009a; Wainwright et al., 2003]) and bacteria that do not form spores (such as micrococcus [Griffin, 2008]). Spore structures allow bacteria to survive for extended periods of time in unfavourable conditions by entering a reversible dormant state. Dormancy is initiated in response to stress [Lennon and Jones, 2011] (Chapter 4), and hence absence of active microbes at a given altitude suggests that the environmental conditions are beyond the tolerance limit for metabolic activity [Nicholson, 2002].
Non-spore forming bacteria that are
retrieved from the atmosphere and are grown in the laboratory, however, may have been active in their environment. Non-spore forming bacteria that are viable but non-culturable have been collected from 20-70 km altitude [Griffin, 2008; Wainwright et al., 2004].
Whether they were metabolically active in their
atmosphere environment is unknown. Both spore forming and non spore forming have difficulty growing at 2.5 kPa [Schuerger and Nicholson, 2005], equivalent to ~ 25 km altitude on Earth [United States Committee on Extension to the Standard Atmosphere, 1976]. There is no confirmed active life at pressures less than 50 kPa (~ 6 km altitude [Gangwar et al., 2009]). The limitations for active life at higher altitudes are not likely to be directly due to pressure. Studies on atmospheric bacteria under martian conditions have found that a small percentage are able to grow at even lower pressures of 350-690 Pa (~ 33-38 km altitude on Earth [Kerney and Schuerger, 2011; Schuerger and Nicholson, 2006]). Although they faced additional stressors of more intense ultra-violet radiation (UV) and an anaerobic atmosphere, the warmer temperatures in the experiments compared to -40 ˚C at the same high altitude pressure environment on Earth likely promoted their growth. Other than temperature and UV radiation [Schuerger et al., 2006; Smith et al., 2010] another strong limiting factor for atmospheric life is likely to be desiccation stress [Diaz and Schulze-Makuch, 2006;
43
Schuerger and Nicholson, 2006]. Water vapour is present in the troposphere and stratosphere [Randel et al., 2004] and it collects as thin films on dust grains and other particles in the atmosphere. Organisms have been found actively growing in liquid water droplets [Amato et al., 2007; Sattler et al., 2001] at temperatures > -15 ˚C. Nutrients are found in clouds and rainwater [Womack et al., 2010] but may be severely limited at higher altitudes above the troposphere. Hence, it is likely that the limitations on nutrients and water availability above ~ 15 km (the average extent of the troposphere) may limit higher altitude active life (rather than low pressure). Furthermore, limitations in temperature may restrict life to ~ 6 km altitude if the low temperature limit of -20 ˚C for life is fundamental [Junge et al., 2006]. It is clear however that low pressure itself has not been shown to be a constraint on active life. Further study is needed to isolate the roles of water, nutrients (and also UV radiation) in constraining atmospheric life [Morris et al., 2011].
2.7.
Temperature
2.7.1. High threshold The highest growth temperature currently known for life is 122 ˚C [Takai et al., 2008]. Hyperthermophiles thrive in environments with temperatures > 80˚C and have physiological adaptations that enable them to live in the high temperature conditions [Amend and Shock, 2001]. Enzymes and proteins in hyperthermophiles remain stable at higher temperature than in other organisms [Gera et al., 2011; Kumar and Nussinov, 2001; Schönheit and Schäfer, 1995], allowing them to maintain structure and metabolism. Most examples of hyperthermophiles are also anaerobic, which is an important adaptation for colonising the deep crust of planets as well as hydrothermal waters [Stetter et al., 1990]. Studying these adaptations is important in determining whether the high temperature threshold for life is higher than the observed maximum. There are a large number of liquid water environments on Earth at higher temperatures (quantified in Chapter 4), which include hydrothermal vents on the ocean floor up to temperatures of 407 ˚C [Koschinsky et al., 2008], subterranean hot springs [Marteinsson et al., 2001], and deep hydrothermal fluids in boreholes [Bucher and Stober, 2010; Pedersen, 2010]. These environments would represent a significant volume of uninhabited liquid
44
Literature Review & Background: Quantifying Habitability & the Extreme Limits of Life
water if life is truly limited by 122 ˚C. Exploration of amino acid denaturation with temperature indicates that life might be fundamentally limited by temperatures above 140 ˚C [Holden, 2008], although other authors hypothesise higher limits of 150 ˚C [Segerer et al., 1993] or even > 200 ˚C [Baross and Deming, 1983; Deming and Baross, 1993; Straube et al., 1990] under the influence of hydrostatic pressure maintaining water as a liquid. A further limitation on life is provided by the increased speed at which chemical reactions progress under high temperature, as if energy is not released in a slow measured way (captured through disequilibrium), life cannot harness it [Gold, 2001; Shock and Holland, 2007].
2.7.2. Low threshold The lowest growth temperature known for life is -20 ˚C [Junge et al., 2006] for psychrophiles (cold loving organisms) isolated from arctic marine sediments. Psychophiles (or psychro-tolerant [Gilichinsky et al., 2007]) have been found in a range of subzero environments growing within liquid brine veins [Streletskaya, 1998] within permafrost and polar ice [Gilichinsky et al., 2003; Junge et al., 2004]. The presence of salts in subzero liquid water environments on the Earth is ubiquitous, but although they enable water to remain liquid at lower temperatures they also result in a lowering of the water activity (2.5) which ultimately can result in desiccation stress and cell death [Mazur, 1980]. Hence life at low temperatures hangs in a delicate balance [Ponder et al., 2008]. Organisms that are active in cold environments maintain a slowed metabolic rate, are able to increase their concentrations of adenosine triphosphate (ATP) and adenosine diphosphate (ADP) to maintain reaction rates in the low temperatures [Amato and Christner, 2009; Deming, 2002; Parry and Shain, 2011], and utilize different enzymes to prevent freezing of their intracellular water [Ayala-del-Río et al., 2010; Laybourn-Parry, 2002]. The low temperature limit of these adaptations is not known [Price and Sowers, 2004], but it is likely that low water activity rather than low temperature provides a more fundamental limit to life in cold liquid water [Gunde-Cimerman et al., 2003]. There are liquid water environments on the Earth colder than -20 ˚C which could provide a viable niche for psychrophiles.
The full range of permafrost
temperatures on Earth can support thin films of liquid coating ice and soil grains
45
[Steven et al., 2006; Tice et al., 1981] with a large fraction able to support brines of NaCl and MgCl2 (Table 2-5). Although solutes have the potential to extend the low temperature limit for life [Chin et al., 2010], they may also exclude life from liquid water if they reduce the water activity to below 0.6 [Grant, 2004]. Table 2-4 reveals than most organisms adapted to low water activity are not from subzero environments. Life in arctic and permafrost environments is generally active in brine with an aw ≥ 0.9 [Ponder et al., 2008], and the most halophilic psychrophiles are found in NaCl brines [Pikuta et al., 2007]. When saturated, an NaCl brine has an activity of 0.76 at its eutectic temperature of -21 ˚C (Table 2-5). Hence, the vast majority of concentrated brines in subzero environments would have an activity > 0.76 and would not remain liquid in temperatures colder than the current low temperature limit of life. As discussed in 2.5.3, psychrophiles are generally well adapted to desiccated environments [Panikov and Sizova, 2007; Panikov et al., 2006]. Considering these arguments, it is plausible that the lack of organisms in both low temperature and low water activity environments (apparent in Table 2-4) is due to the absence of subzero liquid water environments with activity < 0.75. Further studies are needed to identify whether there are natural liquid water environments on Earth below -20 ˚C that have an activity above 0.6, whether these habitats are populated by psychrophiles, and hence if the apparent low temperature limit to life is due to temperature or water activity. Experimental studies to determine if there is positive growth in psychrophiles in -30 ˚C liquid MgNO3 brine [Greenspan, 1977; Morillon and Debeaufort, 1999] could be useful, however organisms that have a high tolerance for salt generally show that adaption only for a particular solute species [KushnerD., 1978].
It is likely
therefore that -20 ˚C is the coldest liquid water environment for active microorganisms on Earth, but that this low temperature cutoff may be different for potential biospheres on other planets with alternative liquid water solutes and hence different subzero liquid water activity.
2.8.
Other factors
There are many other environmental factors which limit the vast majority of organisms, with the exception of a small number of extremophiles that have evolved to tolerate them. For example, at the time of writing this review, active life has been found throughout the pH range of 0-11 [Edwards et al., 2000; Sorokin,
46
Literature Review & Background: Quantifying Habitability & the Extreme Limits of Life
2005; Sturr et al., 1994], and a number of organisms have extremely high tolerance to ultra-violet radiation [Abrevaya et al., 2011], in particular the bacterium deinococcus radiodurans which has no known lethal dosage [Cox et al., 2010] (and is the Guinness Book of Records holder of ‘world’s toughest bacterium’). These thresholds however do not restrict life from a significant fraction of Earth liquid water environments, and hence they have been omitted from this review.
2.8.1. Size of life One final constraint on life in liquid water is the thickness of the liquid water environment relative to the size of the organisms inhabiting it. As the thickness of subzero liquid films found in permafrost and ice is governed by temperature (2.5.2), a size limit for life (or its liquid water environment) is equivalent to a low temperature limit for hospitable liquid water. A possible size limit to life was proposed in response to the ALH-84001, initially believed to house microscopic fossils of extant bacteria [Knoll, 2003].
Nanobacteria remain controversial
[Drancourt et al., 2003], but it is widely proposed that thin films of liquid water are too small for life to grow and metabolise [Ostroumov and Siegert, 1996]. The thickness of these liquid films decreases from ~ 40 Å at 0 ˚C [Anderson and Morgenstern, 1973] to ~ 5 Å (2-3 water molecules) at temperatures below -20 ˚C [Anderson and Tice, 1972; Jakosky et al., 2003]. Films thicker than 20 Å are certainly wide enough for the mobilization of ions [Quinn et al., 2005] and hence provide organisms with a medium for the flow of nutrients and discarding of waste products [Mohlmann, 2009a] and are likely to be habitable [Mohlmann, 2004]. Diffusion barriers within thin films that occur with decreasing temperature [Rivkina et al., 2000] may provide another limit to life, related to low temperature, however the exact threshold is unclear.
47
2.9.
Summary
A number of different metrics are employed by scientists to quantify the habitability of an environment for life with requirements similar to that of Earth’s. In this review the environmental factors that potentially limit life in liquid water environments- temperature, pressure, and water activity- have been discussed. Brief background was also given on two other essential requirements for lifechemical energy and nutrients- which are largely beyond the scope of this thesis. High temperature may constitute a serious limitation for biology, but low temperature is not in itself a serious threat to life. Rather, it is the limitations of water activity that accompanies low temperature that may restrict life from many cold liquid water environments. Also due to their low activity, life may also be restricted from certain brine environments that could be prevalent on other planets. Pressure does not appear to be a severe constraint on life as some organisms thrive in the pressure conditions of the upper atmosphere and the lower regions of the crust. Again, it is the absence of liquid water or low water activity which limits life in these environments on the Earth. Sufficient sources of energy and nutrients may exist on present-day Mars (and more substantially within Mars’ past) to support a subsurface biosphere. The extent and stability of liquid water on present day Mars will be examined in Chapter 5. In this thesis the limitations on microbial life in liquid water to will be applied constrain the astrobiological significance of martian environments and identify the most hospitable regions of the planet for life.
48
Literature Review & Background: Quantifying Habitability & the Extreme Limits of Life
49
3. PHASE DIAGRAM OF EARTH & WATER
‘Consider all this; and then turn to this green, gentle, and most docile earth; consider them both, the sea and the land; and do you not find a strange analogy to something in yourself? For as this appalling ocean surrounds the verdant land, so in the soul of man there lies one insular Tahiti, full of peace and joy, but encompassed by all the horrors of the half known life. God keep thee! Push not off from that isle, thou canst never return!’ -
Herman Melville, ‘Moby Dick’
50
Phase Diagram of Earth & Water
In this chapter the known distribution of liquid water on Earth is quantified and the maximum plausible extent of the hydrosphere predicted. The maximum extent of liquid water within a planet can be quantified through the fraction of environments that have pressure and temperature conditions within the liquid range (2.4). The true extent of liquid water will be less than this upper limit when the actual availability of liquid water is considered, as well as available volume constraints within the subsurface (pore space and permeability). On the surface of Earth and other planets in general, the P-T conditions are determined by the planet’s distance from the sun, the thickness of its atmosphere and the presence of a greenhouse effect. Figure 3-1 illustrates which solar system bodies have the potential for temporary or permanent liquid water on their surface. Within the interior of a planet, the P-T conditions generally increases with depth as the overburden pressure and heat flow increase towards the central core. Hence the interior P-T environment depends on the composition and size of the planet and its heat generation. A pressure-temperature diagram of the Earth is constructed and will be utilised to identify the range of environments where conditions are consistent with the stability of liquid water. Given the necessity of liquid water for life, the maximum depth of liquid water within the Earth is significant as it reflects the deepest possible boundary of the biosphere. Table 3-1: Average pressures and temperatures of some solar system planets Environment
Pressure
Temperature Reference
(Pa)
(˚C)
Earth surface
105
15
[Dziewonski and Anderson, 1981]
Mars surface
600
-50
[Lodders and Fegley, 1998]
Venus surface
9.3×106
480
[Petculescu and Lueptow, 2007]
Titan surface
1.5×106
-180
[Griffith et al., 2000]
Top of Europa’s
1×107
-3.9
[Melosh et al., 2004; Reynolds et
ocean
al., 1983]
51
Figure 3-1: Average surface pressures and temperatures of some solar system planets and moons superimposed on the phase diagram of water. For reference, the average global surface temperatures and pressures for Earth, Mars, Titan, Venus, and at the top of Europa’s potential interior ocean are shown (Table 3-1).
3.1.
Development of the model
3.1.1. Ocean water & largest liquid phase space Details on the phase diagram of pure liquid water and how to adapt the melting and vapourisation curves for aqueous brines (such as seawater) were given in section 2.4. Here those concepts are applied to modelling the phase space of Earth’s liquid water chemistries. Earth’s oceans are dominated by the cations Ca2+, Mg2+, Na+, K+ and the anions Cl-, Br-, SO42-, CO32- and HCO3- [Garrels and Thompson, 1962; Millero, 1969]. A typical ocean salinity is 35 g/L [Duedall and Weyl, 1967] with a freezing point of -2 ˚C at 105 Pa [Fujino et al., 1974] corresponding to ~ 1.16 mol/kg(H2O) of dissolved salts [Dickson and Goyet, 1994]. The corresponding maximum amount of boiling point elevation for sea water is 0.59 ˚C from Equation 2-2. Using this shift of -2 for the melting curve of pure water and +0.59 for the vapourisation curve, the phase diagram of Earth’s ocean water is determined.
52
Phase Diagram of Earth & Water
The full envelope of P-T conditions within the Earth that allow for water, with some fraction of solutes, to remain a liquid corresponds to Earth’s maximum liquid range. Given that liquid water can exist as a thin film to temperatures of at least 100 ˚C [Mazur, 1980; Pearson and Derbyshire, 1974] (2.5.2), but that the coldest temperature environment on the Earth is a surface air temperature of -89 ˚C in Vostok, Antarctica [Ollier, 2007], it is not relevant to include the phase space of colder liquids. Therefore using a freezing point depression of ͺͻ in Equation 2-1
gives a solute molarity of 47.85 moles which provides a boiling point elevation of 25.4 ˚C from Equation 2-2. This shift in the vapour curve of pure water should be viewed only as an upper bound on the true boiling point elevation of a -89 ˚C thin liquid film. The triple point of the largest liquid phase space is at a pressure of ~ 1.3×10-2 Pa and temperature -89 ˚C. The ‘maximum liquid range’ phase space corresponds both to thin films under the full range of temperature conditions that occur on Earth, and to concentrated brines. The extremely sub-zero liquid water (colder than -20 ˚C) would most likely occur only as thin films in icy soil in the polar regions [Davis, 2001; Price, 2000]. Liquid water between -2 ˚C and -20 ˚C occurs both as thin films and brines with a higher molarity than seawater [Duedall and Weyl, 1967; Gilichinsky et al., 2003]. Theoretically, a depression of -89 ˚C to the freezing point of water could be caused by saturated hydrogen chloride salts [Clark and van Hart, 1981], but these salts not common in the natural environment (Table 2-3). The most soluble common inorganic compounds in water have saturated aqueous molarities less than ~ 11 molar [Lide and Haynes, 2009] corresponding to a boiling point elevation of 5.6 ˚C, hence the majority of high and low T values included in the maximum liquid phase space are due to thin films rather than brines.
3.1.2. Earth phase model To examine the full range of pressure and temperature environments on Earth the modelling was divided into sections: atmosphere, surface, oceans, crust, mantle and core. Knowledge on conditions within the atmosphere, surface, oceans and crust is predominantly from direct observation and hence the P-T variability is well constrained.
Within the mantle and core however, pressures and
temperatures are primarily derived from geophysical models, constrained by
53
estimates of densities and mineralogies within the deep interior [Davies, 1999]. Since the early 1900’s it has been recognised that the different rates of propagation of seismic waves travelling through the Earth can provide information on the structure of the interior [Oldham, 1906], and the study of seismology has provided the majority of data used to evolve the understanding of Earth’s deep P-T environment.
It is beyond the scope of this work however to discuss how
geothermal models are developed.
Instead estimates of pressures and
temperatures at depth and their uncertainties from the range of published models will be utilised to develop a new complete model of Earth’s environments. The pressure and temperature conditions within the aforementioned regions of the Earth (atmosphere, surface, oceans, crust, mantle and core) were modelled in two ways.
Firstly, the pressure-temperature boundary of each region was
constructed, giving maximum and minimum values for the plausible P-T conditions within each region. This range in values encompasses two factors – firstly, the true variation in measured conditions within regions of the Earth; and secondly, the uncertainty in modelled P-T conditions arising from the errors associated with the geophysical models used. Once the maximal P-T space of each region of the Earth was constructed, then the representative conditions within each region were illustrated through an average geotherm.
This geotherm indicates the most likely temperature conditions
encountered at a given pressure or depth, determined by the outwards flow of heat from the Earth’s interior and the thermal conduction of the dominant materials. Hence the results based on the Earth phase-space model will be given in the form of mean ± range, where the range includes the uncertainty in geophysical models and the real variation of P-T conditions at a given depth. For the remainder of this chapter as well as chapters 4 and 6, uncertainty refers to model uncertainty, rather than statistical uncertainty. The uncertainty in the boundary conditions in the central core, core-mantle boundary, and crust-mantle boundary (Moho) are represented through vertical and horizontal error bars encompassing the range of estimates in the literature. Within the crust and atmosphere however, the shading is more representative of
54
Phase Diagram of Earth & Water
the range of actual conditions on present-day Earth as there is a more complete and accurate set of P-T measurements available.
3.1.2.a.
Pressure, temperature & depth
The pressure experienced in the Earth’s interior is related to the subsurface depth through:
Equation 3-1
ܲ ൌ ͳͲହ ݃ଵ ߩଵ ݖଵ ݃ଶ ߩଶ ݖଶ ݃ଷ ߩଷ ݖଷ
This pressure is due the overburden of atmosphere (the 1st term) and the layers of crust (2nd term), mantle (3rd term) and core (4th term).
Both gravitational
acceleration g (m/s2) and density ɏ (kg/m3) vary strongly within increasing depth
z (m) within the Earth, and hence a different average value is used for each layer.
The subsurface temperature gradient within the Earth is related to depth through:
Equation 3-2
ௗ் ௗ௭
ൌ
ொ
where the rate of change of temperature with depth
ୢ ୢ
is the geotherm (˚C/m), k is
the thermal conductivity (W/m/˚C) and Q is the heat flow (W/m2).
3.1.3. Atmosphere & oceans The conversion from altitude above the Earth’s surface to pressure, and the geotherm used to plot the range of temperatures experienced with altitude in the atmosphere are given in Table 3-2. Earth’s oceans were represented in the phase model in two ways. Firstly, the most common P-T conditions in the ocean between the surface and the average ocean depth of 3.8 km [Morrissey and Sumich, 2011] were shaded. The minimum ocean temperature is -2 ˚C from the freezing point of seawater at 105 Pa [Fujino et al., 1974], and a reasonable global mean sea-surface temperature is ~ 20 ˚C [Deser et al., 2010; Rayner et al., 2003; Seager et al., 1988]. This gives an estimate of the most common temperature conditions at the Earth’s oceans surface. At the mean depth of the oceans the maximum temperature is 2 ˚C [Fedorov and Ginzburg, 1992]. To model the range of conditions that occur with
55
Earth’s oceans, the data in Table 3-3 was used to draw the ocean boundary polygon, encompassing the full variation of pressure and temperature across the 11 km maximum ocean depth [Kato et al., 1998]. This ocean region includes hydrothermal vents and supercritical water. Table 3-2: Model parameters for Earth’s atmosphere. Parameter
Value
Mean temperature lapse
6.5 troposphere
rate (K/km)*
-2.0 stratosphere 2.6 mesosphere -5.8 thermosphere
Pressure gradient (Pa/km)
4.26
** (below 86 km)
z æ ö - 0.0012 ´ ç1 ÷ 44307 . 69 è ø
References
[Choi et al., 1998; Hewitt and Jackson, 2003; Pearson and Rugge, 1976; United States Committee on Extension to the Standard Atmosphere, 1976; West, 1996]
*Temperature lapse rate is the rate of decrease of temperature with height **Pressure gradient is the rate of decrease of pressure with height dP/dz using the equation: z = [ 1 - (P/10.13)0.19 ] x 44307.69 where z = altitude (km) and P = pressure (Pa) [Pearson and Rugge, 1976]
3.1.4. Surface & crust 3.1.4.a.
Boundary
Extreme environments of the crust and surface of Earth are given in Table 3-3 and are used to define the pressure-temperature boundary of these regions in the Earth P-T model.
Each of the environments in Table 3-3 has been directly
sampled/measured. As they represent extreme P-T conditions, all environments within the crust and on the surface lie within the pressure-temperature polygon defined by these points, labelled ‘crust’ in Table 3-4 (the term crust now implicitly includes the surface region). The crust boundary environments can be described by the phrases (with reference to Table 3-3): hot, low pressure (A); hot surface (B); hot, high pressure (L,M,N,O); cold, low pressure (C,D); cold surface (E), cold, high pressure (F,G,H). The remaining letters in Table 3-3 describe the crust-mantle boundary at the base of the oceanic and continental crust, and the upper mantle (3.1.5). It is assumed that all the points within the crust polygon exist on the Earth.
56
Phase Diagram of Earth & Water
This is a reasonable assumption given that the Earth is not a closed system and hence heat flow and mass flow occur between neighbouring environments. To represent the mean conditions within the crust and hence the most likely P and T conditions that would be encountered at a given depth, the variation in crustal thickness, heat flow, and geothermal gradients is examined below. Table 3-3: Extreme temperatures and pressures in Earth’s oceans and crust Key
Environment
A
Surface lava and hottest surface temperatures (Cotopaxi, Ecuador)
B
Thermal spring/vent (Pork Chop Geyser, Yellowstone, USA) Mt Everest summit
C,D E
Minimum land temperature (Vostok, Antarctica) Deep ice (South Pole, Antarctica) Deep ice (Vostok, Antarctica)
F G H, K I
J
L M
N O
Marianas trench (Challenger Deep) Mean base of oceanic crust
Crust mantle boundary – lowest temperature at thickest point (Himalayas, southern Tibet) Sea floor hydrothermal vents – hottest (Mid-Atlantic ridge) Sea floor hydrothermal vents – hottest shallow vents (Tonga arc) Geothermal waters in the crust Lava lakes (Kilauea Iki, Hawaii; Eldfell volcano, Iceland)
Pressure (Pa) Depth (km)* 4×104 -5.897
Temp (˚C) 1200
7.7×104 -2.293 (3.3 ± 0.1)×104 -8.848 6.0×104 -3.488 7.2×106 0.8 2.5×107 2.75 1.32×108 10.896 2.0×108 10
97
5×109 90
Reference
-74 to -26
[Decker and Decker, 1992; Fagents and Greeley, 2001; Harris et al., 1998; Stein and Spera, 2002] [Guidry and Chafetz, 2003] [West et al., 1983]
-89
[Ollier, 2007]
-51
[Price et al., 2002]
-40
[Abyzov et al., 1998]
2 to 300
500
[Kato et al., 1998; Todo et al., 2005] [Huppert and Sparks, 1988; Lodders and Fegley, 1998] [Priestley et al., 2008]
3.5×107 3 4.9×106 0.54
464
[Koschinsky et al., 2008]
265
[Stoffers et al., 2006]
2.8×106 0.1 107 0.3
157
[Ellis and Mahon, 1977]
1500
[Helz and Thornber, 1987; Jonsson and Matthiasson, 1974]
200
*Conversions between depth and pressure for atmosphere; land and water used the density values in Table 3-4, Table 3-2 and Equation 3-1. Gravitational acceleration values were taken from [Dziewonski and Anderson, 1981].
57
3.1.4.b.
Average
The boundary between the crust and the mantle beneath is marked by a density transition known as the Mohorovicic discontinuity (Moho), discovered in 1909, [Mohorovicic, 1909], where seismic wave velocities increase with depth. It does not have a fixed pressure or temperature although the global average depth for the Moho is 19 km [Dziewonski and Anderson, 1981]. The range of depths for the transition from crust to mantle varies widely. The thickness of continental crust ranges from 5-90 km (averaging 35 km [Holbig and Grove, 2008; Pavlenkova and Zverev, 1981]) but varies between only between 2-10 km for oceanic crust (averaging 7 km [White et al., 1992]). The oceanic crust is relatively uniform, except at mid-ocean ridges where the mantle is exposed and new crust is generated. In contrast the continental crust is much more variable, with the thickest portion underlying the mountains of the Himalayas in the Tibetan plateau [Wittlinger et al., 2004] and the thinnest crust occurring in continental rift zones [Thybo and Nielsen, 2009]. Considering first the continental crust: a reasonable average temperature for the Moho beneath mean ~ 35 km thick crust is 800 K (527 ˚C) [Blackwell, 1971; Mooney et al., 1998], within a range of ~ 573-1400 K due to the heterogeneity of heat flow and geothermal gradients [Hyndman et al., 2005] (discussed further in 3.1.4.c below). Taking an average density of 2800 kg/m3 [Lodders and Fegley, 1998] provides a pressure of 9.6×108 Pa at the base of the crust from Equation 3-1. The representative pressures at the thinnest 5 km crust and thickest 90 km crust can also be calculated, and provide the range given in Table 3-4 [Sokolova et al., 1978]. The temperatures within the crust generally increase with depth and hence the thickest crust is expected to be hottest. There are however regions of deep cool crust and Moho which contradict this trend. These typically occur beneath cratons- ancient and immobile regions of the crust that have not undergone deformation for ~ 104 Ma [Gao et al., 2002; Levy and Jaupart, 2011]. Considering the variation of the Moho beneath the thickest crust, the temperatures from current models are between 800-1200 K [Jiménez-Munt et al., 2008; Priestley et al., 2008].
58
Phase Diagram of Earth & Water
Table 3-4: Model parameters for Earth’s crust. Continental crust Thickness (km) Temperature (K)
Mean
Range (incorporating measured variation and model uncertainty) 5-90 Surface: 184 - 331 Base: 573 - 1400 Surface: 1.1×105 – 3.3×104 Base: 1.4×108 – 2.5×109 N/A N/A N/A
35 Surface: 288 Base: 800 Pressure (Pa) Surface: 105 Base: 9.6×108 Density (kg/m) 2.8×103 Gravity (m/s2) 9.8 Pressure gradient 2.74×104 (Pa/km) Geothermal gradient 25 5-70 (K/km) References [Blackwell, 1971; Burtman and Molnar, 1993; Cermak et al., 1989; Dziewonski and Anderson, 1981; Holbig and Grove, 2008; Huppert and Sparks, 1988; Hyndman et al., 2005; Jiménez-Munt et al., 2008; Lodders and Fegley, 1998; Mooney et al., 1998; Pavlenkova and Zverev, 1981; Pollack and Chapman, 1977; Scheidegger, 1976; Sokolova et al., 1978; Thybo and Nielsen, 2009; Wittlinger et al., 2004] Oceanic crust Mean Range Depth (km) 3.8 0-11 Thickness (km) 5 4-20 Temperature (K) Surface: 274 Surface: 271 – 737 Base: 473 *intersects with continental crust geotherm at 473 Pressure (Pa) Surface: 3.8×107 Surface: 105 - 1.3×108 Base: 1.6×108 – 6.2×108 Base: 2×108 *intersects with continental crust geotherm at 2×108 Density (kg/m3) 3.0×103 N/A 2 Gravity (m/s ) 9.8 N/A 4 Pressure gradient 2.94×10 N/A (Pa/km) Geothermal gradient 35 N/A (K/km) References [Anderson, 1979; Chapman, 1986; Davies, 1999; Dziewonski and Anderson, 1981; Huppert and Sparks, 1988; Koschinsky et al., 2008; Lodders and Fegley, 1998; Mooney et al., 1998; Muller and Smith, 1993; Scheidegger, 1976; Sclater and Francheteau, 1970; Sclater et al., 1980; White et al., 1992]
59
It should be noted that additional variation in lithostatic pressure occurs due to the density variation of rocks within the continental crust, with typical densities of andesitic rocks ranging from 2600-3250 kg/m3 across the crust, increasing with depth [Herzberg et al., 1983; Rudnick and Fountain, 1995]. Hence, using a constant density of 2800 kg/m3 for the continental crust could introduce an error of at most 17% in the pressure calculation. This is not significant, as although it would alter the modelled representative pressure of Moho, the estimate would remain within the range of uncertainty in geophysical models (Table 3-4). An additional 1-3% of variation in the crustal density is introduced in the thickest and coolest portions of the crust (cratons) [Kaban et al., 2003]. The methodology for representing oceanic crust was identical and used values and references given in Table 3-4. The mean P-T conditions within the crust and the Moho are represented by an average geotherm.
Error bars encompass the
variation of typical P-T conditions at the base of the crust. This variation in pressure arises due to the heterogeneity of crustal thicknesses and density (discussed above). The temperature variation occurs due to the range of crustal conductivity and heat flow, discussed below.
3.1.4.c.
Geothermal gradients in the crust
A fundamental parameter in the model which is not explicitly represented is Earth’s heat flux. The heat flux in combination with the thermal conductivity of the subsurface establishes the rate at which temperature increases with depth (the geotherm, Equation 3-2). The heat flow measured at the surface of the earth combines a ~ 40% contribution from radioactive decay within the crust and ~ 60% from the heat production of the core [Pollack and Chapman, 1977], although this varies strongly with location [Chad-Umoren and Osegbowa, 2011]. Heat flux has a complex relationship with depth, as radioactive heat production varies with rock type, density and pressure [Gordienko and Pavlenkova, 1985; Rybach and Buntebarth, 1984]. Earth’s surface heat flow varies predominantly between 40-85 mW/m2 [Pollack and Chapman, 1977] with an average of ~ 55 mW/m2 (Figure 3-2). The heat flow from the Moho beneath the continents (average ~ 35 km depth) is fairly fixed at around 25 mW/m2 [Cermak et al., 1989] with the additional surface heat flow from radiogenic heat production within the crust. Heat flow
60
Phase Diagram of Earth & Water
from the Moho beneath oceanic crust is much higher with an average of 100 mW/m2 [Davies, 1999; Sclater and Francheteau, 1970]. There is less radiogenic heat production in the oceanic crust itself [Sclater et al., 1980] and thus its heat flow shows much less variation.
Figure 3-2: Geothermal heat flow map of north America2. Range here is 25-29 W/m2 (light blue bin) to 100-150 mW/m2 (dark red bins) [Blackwell and Richards, 2004]. The map reveals the heterogeneity in heat flow which results in significant variation in geothermal gradients within the crust, and hence in the subsurface temperature environment. The heterogeneity in heat flow and the conductivities of crustal materials (for example, a range of 0.5–7W/m/K for rock [Clauser and Huenges, 1995]) has the consequence that temperatures can still show large regional variation many kilometres beneath the surface [Artemieva, 2006], even continuing into the asthenosphere [Smith, 1973] (the ductile region of the mantle). Thus oceanic and continental geotherms cannot be guaranteed convergence until below ~ 200 km in the asthenosphere where convective heat transfer dominates [Anderson, 1979; Chapman, 1986]. In the P-T model, the average oceanic and continental geotherms converge at ~ 10 km into the continental crust and at the base of the thickest 2
http://smu.edu/geothermal/heatflow/heatflow.htm
61
oceanic crust (point I, Table 3-3), however the range of shading in the crust is also consistent with deeper convergence. Using Equation 3-2, a typical geothermal gradient in the oceanic crust is ~ 35 K/km, corresponding to a 100 mW/m2 heat flow, a thermal conductivity for basalt of 3 W/m/K [Clauser and Huenges, 1995]. The mean continental geothermal gradient in the continental crust is 25 K/km [Scheidegger, 1976] but ranges between 5-70 K/km [Chapman and Pollack, 1975; Pollack and Chapman, 1977]. In the P-T model, the mean crustal geotherm, averaged over 35 km of crust, is 15 K/km.
The crust region (to 90 km beneath the surface) accommodates the
aforementioned range of geotherms in addition to steeper gradients up to ~ 80 K/km, and so the temperatures within the crust are consistent with the range of true geothermal gradients on Earth. The temperature range of ~ 573-1400 K for the Moho beneath 35 km crust is consistent with geothermal gradients between 6 K/km (relevant to low heat flow regions, such as from point H to the average base of the crust) and 38 K/km (from point E to the average base of the crust).
3.1.5. Mantle & core Earth’s mantle extends from depths ranging between 4-90 km beneath oceanic and continental crust to 2890 km at the core-mantle boundary (CMB).
A
representative pressure value for the CMB of 1.3×1011 Pa is used in the model [Knittle and Jeanloz, 1991], within a range of (0.6-3.2)×1011 Pa [Davies, 1999; van der Hilst et al., 2007; Lay et al., 1998]. The representative pressure is consistent with using a depth of 2890 km from the surface, under continental crust 35 km thick, density and gravity parameters given in Table 3-5, and pressure calculated from Equation 3-1. The representative CMB temperature is 2900 K (2627 ˚C) [Kamada et al., 2010; Lay et al., 2008], within a range of 1900-4500 K [Davies, 1999; van der Hilst et al., 2007; Jeanloz, 1990; Knittle and Jeanloz, 1991]. It is important to note that although it is represented in the P-T model as occurring at a fixed depth, the transition from core to mantle occurs within a thermal boundary layer ~ 40 km thick [Doornbos, 1983; Williams and Garnero, 1996] with an estimated radial temperature increase of 800 ˚C across it [Stacey and Loper, 1984]. The modelled CMB boundary values lead to a mantle geotherm of 0.66 K/km. This temperature profile is cooler than the temperature profile of a liquid mantle
62
Phase Diagram of Earth & Water
[Nakagawa and Tackley, 2004, 2010; Zerr et al., 1998], hence resulting in the expected solid mantle [Stein and Wysession, 2003].
For reference, the CMB
temperature (with the corresponding representative CMB pressure from Table 3-5) which would result in a liquid mantle is estimated to be 4300 K, which is 1400 K hotter than the CMB in this model [Nakagawa and Tackley, 2004]. Earth’s core extends from a depth of 2890 km to 6378 km [Davies, 1999]. In the centre the pressure is modelled as 3.6×1011 Pa [Dziewonski and Anderson, 1981; Sata et al., 2010; Stixrude and Cohen, 1995]. As the core density is not accurately known from seismic models, there is a range of estimates for the central core pressure between (2.8–3.8)×1011 Pa [Asanuma et al., 2011; Fowler, 2005; Kuwayama et al., 2008; Lodders and Fegley, 1998]. At its central depth, the core is modelled as having a representative temperature of 5700 K (5427 ˚C) [Fowler, 2005; Sata et al., 2010]. Temperature models for the core are dependent on poorly-constrained core composition and hence estimates for the central temperature range widely from 3800-6800 K [Jeanloz, 1990; Saxena et al., 1994; Stixrude et al., 1997]. The core boundary values from the P-T model lead to a geotherm of 0.85 K/km across the core. Melting temperatures for the assumed core composition are pressure dependent [Anderson, 2002], however the geotherm in this model results in a liquid outer region of the core and a solid inner region (from the melting profile of [Jeanloz, 1990; Nakagawa and Tackley, 2004]). This is consistent with recent models [Stein and Wysession, 2003].
63
Table 3-5: Model parameters for Earth’s core and mantle. Mantle
Mean
Depth (km)
Outer: 35 Base: 2890 Outer: 800 Base: 2900 Outer: 9.6×108 Base: 1.3×1011 4.5×103 10.3 4.4×104
Temperature (K) Pressure (Pa) Density (kg/m3) Gravity (m/s2) Pressure gradient (Pa/K) Geothermal gradient (K/km) References
Core Depth (km) Temperature (K) Pressure (Pa) Density (kg/m3) Gravity (m/s2) Pressure gradient (Pa/K) Geothermal gradient (K/km) References
0.7
Range (incorporating measured variation and model uncertainty) Outer: 5 - 90 Outer: 573 - 1400 Base: 1900 - 4500 Outer: 1.4×108 – 2.5×109 Base: (0.6 - 3.2)×1011 N/A 9.9 - 10.5 N/A 0.3 - 1
[Davies, 1999; Decker and Decker, 1992; Doornbos, 1983; Dziewonski and Anderson, 1981; Fowler, 2005; Garnero and McNamara, 2008; Garnero, 2004; van der Hilst et al., 2007; Jeanloz, 1990; Kamada et al., 2010; Knittle and Jeanloz, 1991; Lay et al., 1998, 2008; Nakagawa and Tackley, 2004, 2010; Scheidegger, 1976; Schiano et al., 2006; Stacey and Loper, 1984; Stein and Wysession, 2003; Williams and Garnero, 1996; Zerr et al., 1998] Mean Range Outer: 2890 N/A Centre: 6378 Outer: 2900 Outer: 1900 - 4500 Centre: 6100 Centre: 3800 - 6800 Outer: 1.3×1011 Outer: (0.6 - 3.2)×1011 11 Centre: 3.6×10 Centre: (2.8 - 3.8)×1011 1.3×104 N/A 5.5 0 - 10.5 5 1.3×10 N/A 0.85
0.6-0.9
[Anderson, 2002; Asanuma et al., 2011; Davies, 1999; Doornbos, 1983; Dziewonski and Anderson, 1981; Fowler, 2005; van der Hilst et al., 2007; Jeanloz, 1990; Kamada et al., 2010; Knittle and Jeanloz, 1991; Kuwayama et al., 2008; Lay et al., 1998, 2008; Lodders and Fegley, 1998; Nakagawa and Tackley, 2004; Sata et al., 2010; Saxena et al., 1994; Schiano et al., 2006; Stacey and Loper, 1984; Stein and Wysession, 2003; Stixrude and Cohen, 1995; Stixrude et al., 1997; Williams and Garnero, 1996]
64
Phase Diagram of Earth & Water
3.1.6. Transient The dynamic nature of the Earth is represented in part through the region labelled ‘transient’. Pressure and temperature conditions within this region indicate shortterm environments that are temporarily at typical mantle and deep crustal temperatures but are at crustal pressures. They include magma chambers in the crust, surface lava, and subsurface lava lakes [Helz and Thornber, 1987; Jonsson and Matthiasson, 1974].
Over time-scales of decades, the majority of these
environments will cool along a constant pressure line to lie within the crustal region [Oppenheimer and Francis, 1998].
The abundance of shallow, hot
environments described by the transient region of phase space would be expected to decrease over time as the mantle continues to lose heat at the rate of ~ 100 ˚C per million years [Tackley, 2000]. The extreme boundary of the transient region occurs at the lowest pressure and highest temperature environment on the Earth, corresponding to the hottest lava at the highest altitude volcano (point A, Table 3-3). On the right hand side beneath point A, the transient region is bounded by a geothermal gradient of ~ 0.6 K/km. This is an appropriate geotherm for deep mantle material (3.1.5). Much steeper geotherms appropriate for shallow lava chambers [Armienti et al., 1984; Hardee and Larson, 1977] are also accommodated within the transient region, for example a geothermal gradient of 200 K/km beneath point B (Table 3-3).
3.1.7. Interpreting the Earth phase diagram 3.1.7.a.
Temporal variation
One limitation of representing all the Earth’s environments on a static twodimensional model is that it does not allow the temporal variability of pressure and temperature conditions to be explicitly illustrated. For a given P-T value, there is limited information in the model on the probability of such an environment occurring within the Earth, or how long such conditions are maintained. The Earth phase diagram indicates that a given P-T value has either been measured in one or more locations in the Earth, or has a high probability of occurring from the predictions of geothermal models. The average geotherm does however provide an estimate of the most likely configuration of pressure and temperature at a given depth as it is based on mean heat flow, thermal conductivity and density. Despite
65
not being explicitly represented, the model is consistent with the temporal variation in geotherms. A few examples are given. The transient region (3.1.6) reflects short-term hot environments within the Earth which will eventually cool to lie within the crust region. Another example of temporal variation in pressure and temperature incorporate in the model is the hydrological cycle of the Earth. The phase changes of water drive the exchange of water between surface and atmosphere (evaporation, condensation, sublimation), through the variation in the partial pressure of water in the atmosphere. This dynamic nature is implicitly incorporated in the P-T model as the width of the phase space encompasses the cycling of pressures and temperatures at a given depth/altitude. For example, the temperature at the summit of Mount Everest varies between the maximum (point C) and minimum (point D), and atmospheric pressures and temperatures vary within the pink shaded region around the average atmospheric geotherm.
3.1.7.b.
The meaning of pressure
The vertical pressure axis on the Earth and liquid phase model holds a dual meaning. For the phase changes of water, the relevant pressure is the partial pressure of water vapour (2.4).
For the Earth’s environments, the relevant
pressure is the force exerted by an overburdening mass on a unit area of surface (Equation 3-1). On the surface and within the atmosphere the pressure is supplied by the overburden of gas (average of 105 Pa at sea level). Within the ocean hydrostatic pressure corresponds to the overburden of water; and within the subsurface lithostatic pressure corresponds to the overburden of rock.
The
difference between the meaning of the vertical pressure axis for water and for the Earth however could lead to a difficulty in interpreting the pressure of water environments within the subsurface and what phase of water is dominant. The relationship between the pressure of subsurface liquid water within pore space and depth is complex, as it depends on how the liquid is confined and whether it is in contact with an environment of lower partial pressure. For example, if subsurface pore space water at a depth of 500 m is in vapour contact with the atmosphere, the pressure exerted on the surface of the liquid will be 105
66
Phase Diagram of Earth & Water
Pa (or ~103 Pa of water vapour)- significantly less than the ~ 107 Pa lithostatic pressure under 500 m of rock. The closure depth of pore space varies widely with substrate and grain size, but can generally be assumed to be effectively closed to the atmosphere beneath ~10’s of centimeters [Heard et al., 1988]. There are many examples of confined aquifers on Earth 1–3 km deep where the water is not connected to the surface atmosphere, and hence the pressure on the water is given by or exceeds hydrostatic pressure [Orr and Kreitler, 1985]. Overburden pressure is used on the vertical axis in the Earth phase model and hence any liquid water within the pore space of rock is indicated as being at the lithostatic pressure experienced by the adjacent rock. The density of rock in both the crust and upper mantle is at most three times the density of water [Dziewonski and Anderson, 1981].
This factor of three between hydrostatic pressure and
lithostatic pressure has a minimal effect on the phase space occupied by the Earth, given the range of pressures encompassed by the error bars at the Moho and coremantle boundary. Hence it is a reasonable estimate to use lithostatic pressure for confined subsurface liquid water environments within pore space beneath the closure depth.
Within the ocean hydrostatic pressure is naturally used, as
indicated by the ‘ocean depth’ axis.
3.2.
Results: the pressure-temperature representation of Earth &
liquid water Figure 3-3 presents the phase space of liquid water chemistries relevant to the Earth. The dotted curve marks the phase space boundary of ‘ocean water’. The solid curves mark the modelled ‘maximum liquid range’- the full extent of phase space occupied by liquid water (fresh water, salt water, concentrated brines, thin films, and super-heated water) on Earth.
The phase diagram of all Earth
environments is shown in Figure 3-4, superimposed on the phase diagrams of water. Conceptually, this shows the regions of phase space where there is both Earth and liquid water, where there could be liquid water but there is no Earth, and where there is Earth but no liquid water. The first and the last are most useful in investigating the maximum extent of habitable environments on Earth, which will be done in Chapter 4. The second may be important for Mars (Chapter 6).
67
Figure 3-3: Phase diagram of water chemistries found on Earth. Triple point of the ‘maximum liquid range’ is -89 ˚C (184 K) and 0.013 Pa. Triple point of ‘ocean water’ is -2 ˚C (271 K) and 600 Pa. The bulk of Earth’s ocean water is shown in the narrow dark blue wedge.
3.2.1. Implications of high solute & thin film liquids The narrow wedge in Figure 3-3, indicating the P-T conditions within the bulk of Earth’s oceans, demonstrates that only a small fraction of liquid water environments on the Earth will occur within the remainder of the liquid phase space. Although the subzero liquid region is appropriate for the concentrated brines and thin liquid films within ice, permafrost and coating atmospheric grains on Earth, most low temperature water environments have only a moderate freezing point depression. The freezing points of the most common fluids in cold environments on the Earth - Na+, Mg2+ and Ca2+ chloride brines are around -10 ˚C [Duedall and Weyl, 1967] (2.4.2). From Figure 3-4, liquid water with typical ocean salinity can exist over the majority of the Earth’s surface at some stage during seasonal cycling. It is restricted only from the highest altitudes and coldest surface conditions (such as points C and D in Table 3-3). As the phase diagram of thin films (and concentrated brines) has a triple point at -89 ˚C and 1.3×10-2 Pa, it
68
Phase Diagram of Earth & Water
represents liquid which can occur throughout the full range of Earth’s surface conditions and up to an altitude of ~ 86 km within the atmosphere. Within the crust region of Figure 3-4 the average geotherm is always within the liquid regime of ocean water. Much lower subsurface temperatures however occur in the nearpolar environments. The thermal conductivity of polar ice is typically 2 W/m/K [Pollard et al., 2005], with a heat flow around 40 mW/m2 [Ritz et al., 2001]. Using Equation 3-2 this provides a typical polar geotherm of ~20 K/km [Budd and Jenssen, 1989]. Hence, using the annual average surface temperature of -50 ˚C in Antarctica3, the average depth of the -2 ˚C melting isotherm of ocean salinity water is 2.4 km (for comparison, Lake Vostok is ~ 4 km beneath the surface of the ice sheet [Siegert et al., 2001]). Therefore from the P-T space thin films would be expected to be the dominant form of crustal liquid water in cold regions of the Earth, down to a depth of at least 2 km (and extending deeper in regions where the substrate has a higher thermal conductivity such as point G, Table 3-3). This is consistent with observations of polar soils, as thin films are the primary source of liquid water wherever the annual average soil temperature is ≤ -10 ˚C [New et al., 2002]. On the next page: Figure 3-4: Earth phase diagram superimposed on that of liquid water, with average geotherm and error bars at boundary regions. The second geotherm between ~ 4-10 km ocean depth corresponds to the geotherm through the oceanic crust. The blue vertical wedge identifies the majority of ocean water and the blue shading indicates the range of all liquid water environments on Earth (from Figure 3-3).
3
http://www.antarcticconnection.com/antarctic/weather/climate.shtml
Figure 3-4
69
70
Phase Diagram of Earth & Water
3.2.2. Deepest liquid water The model prediction of the deepest liquid water on Earth is shown by the red asterisk in Figure 3-5 (and Table 3-6). It corresponds to the intersection of the maximum pressure of the Earth phase space with the critical point temperature of ocean liquid water. This environment occurs in the crust at 374 ˚C (647 K) and 2.5×109 Pa, corresponding to a subsurface depth of ~ 70 km and hence generally lying within the mantle. As this is the maximum extent of liquid water on Earth it represents the deepest possible boundary of the biosphere. Liquid water is stable at higher pressures than this limit but Earth environments at these pressures do not exist. For example, ocean water exists as a liquid at 300 ˚C and 3×109 Pa but the Earth’s phase space does not include these conditions. The maximum extent of liquid water within the Earth corresponds to the hydrosphere extending across at most ~ 3.3 % of the total volume of the Earth.
Figure 3-5: Simplified plot of Earth’s phase space and the phase space of Earth water. The blue shading indicates liquid range of water on Earth.
The red asterisk
represents the model prediction of the hottest and deepest possible water on Earth.
71
As the phase space of the Earth is modelled from geophysical estimates and uncertainties of pressure and temperature gradients with depth, it is possible that the boundaries of the Earth’s crust and mantle will change as models improve. This may change the aforementioned estimate of the deepest liquid water. The robustness of this estimate can be assessed by examining the range of geothermal gradients that are consistent with the Earth’s modelled P-T space. Figure 3-4 represents the average P-T value for the crust-mantle boundary as a diamond with error bars reflecting the variation and uncertainty in temperature, depth and pressure of this boundary. The shallowest Moho beneath continental crust occurs at a depth of 5 km and at temperatures between ~ 200-500 ˚C [Blackwell, 1971; Huppert and Sparks, 1988; Mooney et al., 1998]. Deep Moho beneath thick crust of 90 km occurs in the temperature range of 500-1200 ˚C [Hyndman et al., 2005; Jiménez-Munt et al., 2008; Priestley et al., 2008].
This range of crust-mantle
boundary conditions requires geotherms between 6-49 K/km, which is consistent with the model. Unless the thickest regions of the Earth's crust are found to exist above mantle regions with temperatures below 200 ˚C, the maximum depth for liquid water predicted by this model will be 70 km. Table 3-6: Pressure and depths of features of interest Description
Pressure
Temp. Pressure-depth
Depth
(Pa)
(˚C)
constants*
(km)
374
h1 = 3.5×104
14.5
Hottest liquid water along 3.98×108 mean geotherm Hottest and deepest possible 2.51×109
ρ1 = 2800 374
h1 = 3.5×104
liquid water (red asterisk in
ρ1 = 2800
Figure 3-5)
h2 = 3.5×104
70.0
ρ2 = 4500 Closure of pore space along 3.16×108
350
mean geotherm Maximum depth estimate of 2.24×109 pore space closure
h1 = 3.5×103 ρ1 = 2800
350
h1 = 3.5×103 ρ1 = 2800 h2 = 2.8×104 ρ2 = 4500
*Relevant to Equation 3-1
11.5 63.4
72
3.2.2.a.
Phase Diagram of Earth & Water
Pore space & fluid access
Despite P-T conditions that allow for water to be in its liquid phase at 70 km depth (at 374 ˚C), it is possible that water would be restricted from such an environment by a lack of pore space. Pore space declines exponentially with increasing depth due to the increasing overburden pressure [Athy, 1930]. The measured crustal porosity is generally less than 5 % beneath ~ 5 km depth in continental crust [Pechnig et al., 1997] and beneath ~ 1 km in oceanic [Becker, 1985].
The
percentage porosity at a depth z (m) can be modelled by:
Equation 3-3
ሺݖሻ ൌ ݁ ି௭
where is the porosity at a reference depth and the rate of decrease depends on
the compaction coefficient k (m-1) [Athy, 1930]. In practice however, a wide range of porosities are encountered at a given depth but the maximum porosity encountered consistently decreases with increasing depth according to the theoretical relation in Equation 3-3 [Jamialahmadi and Javadpour, 2000]. Liquid water is found in rocks with extremely low porosity, as fractures filled with brine have been found in continental rock at 11 km (Soviet ultradeep Kola SG3 borehole [Jodicke, 1992]) while estimated porosities for crustal rocks at this depth range from 0.5-3 % [Hyndman and Shearer, 1989]. Porosity at pressures above 100 MPa is generally assumed to be zero [Christensen, 1973]. Despite this, modelling of crustal rocks at pressures of up to 350 MPa under a range of temperatures, has revealed porosities around 1 % for amphibole, granulite and pyroxene granulite rocks [Abdulagatov et al., 2006]. Rather than by pressure, pore fluid may be more strongly limited by temperature, existing only to a limiting temperature of ~ 350 ˚C [Hyndman and Shearer, 1989]. This temperature is consistent with 3.16×108 Pa and 12 km depth within the Earth along the average geotherm (Figure 3-4). A temperature limit for pore space is not due to the phase space of water but rather to the extreme deformation of pore space in all rocks at this temperature which restricts the available volume for water. Considering therefore that no rocks that have experienced temperatures of > 350 ˚C are expected to have significant pore space due to temperature induced deformation, it is plausible that liquid water may be restricted to depths above ~ 63 km (from the maximum depth of 350 ˚C in the Earth phase space, Table 3-6).
73
3.2.3. Liquid water budget The amount of liquid water on Earth that is present in the deep interior can be estimated. This is distinct to the volume of Earth within the hydrosphere, which is the fraction of Earth’s volume that can support liquid water. The total volume of water on Earth is ~ 1.39×109 km3 of which ~ 98% or 1.36×109 km3 exists as liquid water [Gleick, 1993; Gleick et al., 2008]. There are a range of estimates for the total volume of subsurface liquid water in the crust, from (0.7-2.3)×107 km3 [Foster and Chilton, 2003; Lehr and Lehr, 2000; Rogers and Feiss, 1998], dependent on estimates of subsurface fracturing, porosity and groundwater depletion [Konikow and Kendy, 2005]. An estimate for the volume of groundwater shallower than 4 km is 9.5×106 km3 [Whitman et al., 1998], at which depth the mean temperature is ~ 120 ˚C (from the average geotherm in Figure 3-4). Therefore by taking the ratio with the total volume of liquid water within the crust indicates that > 41.3 % of ground water is estimated to be shallower than 4 km depth. Hence only a very small amount of the Earth’s liquid water budget (< 107 km3) potentially extends to the maximum depth of liquid water stability at 70 km. This volume of liquid water is also an estimate of the amount of liquid water on Earth that is too hot for life, as the current highest temperature life in liquid water is found at 122 ˚C [Takai et al., 2008] (2.7.1). The estimate of subsurface groundwater beneath 4 km ignores the < 5 wt% of water in the upper mantle estimated to comprise ~10-9 oceans [Litasov and Ohtani, 2005; Ohtani, 2005] (volume varying with mineral type and burial depth). This water exists as OH within hydrous minerals rather than free liquid H2O. To date, the deepest groundwater that has been sampled is at 11 km depth, filling open rock fractures within the Kola SG3 borehole [Jodicke, 1992]. There is also a substantial volume of thin film liquid water within permafrost (2.7.2). The amount of unfrozen water in soil depends strongly on the soil content and vapour pressure of water, but also on the fraction of ice- with the amount of liquid film increasing with increasing ice content [Nakano et al., 1982]. Typical values of liquid water in frozen soil are 3.5-15 % at -25 ˚C [Cary and Mayland, 1972; Dillon and Andersland, 1966; Yoshikawa and Overduin, 2005].
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3.3.
Phase Diagram of Earth & Water
Summary
In this chapter a phase model for the Earth and its liquid water was constructed, by representing the full extent of pressure and temperatures conditions within the environments. Earth’s hydrosphere can extend to at most ~ 70 km beneath the surface, encompassing at most 3.3 % of the planet’s total volume. The work presented here provides a new comprehensive synthesis on current knowledge of Earth’s pressure and temperature environments, which is then applied to quantify the mean and maximal extent of liquid water within the Earth through detailed depth and volume estimates. This model will be used to provide new information on the biosphere in Chapter 4, through comparing the range of pressures and temperatures of environments that support life. This will allow the fraction of hospitable liquid water environments on the Earth to be quantified. In Chapter 6 the phase space of Mars and martian water will be modelled and the putative extent of the martian hydrosphere will be compared to Earth’s.
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4. PHASE DIAGRAM OF EARTH’S BIOSPHERE
‘And while they walked Jolyon pondered those words: ‘I hope you’ll treat him as you treated me.’ That would depend on himself. Could he trust himself? Did Nature permit a Forsyte not to make a slave of what he adored? Could beauty be confided to him? Or should she not be just a visitor, coming when she would, possessed for moments which passed, to return only at her own choosing? ‘We are a breed of spoilers!’ thought Jolyon, ‘close and greedy; the bloom of life is not safe with us. Let her come to me as she will, when she will, not at all if she will not. Let me be just her stand-by, her perching-place; never- never her cage!’’ -
John Galsworthy, ‘The Forsyte Saga: In Chancery’
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Phase Diagram of Earth’s Biosphere
All known life on Earth requires liquid water.
Although it is plausible that
organisms on other planets may have unique requirements and utilise biochemistry vastly different from Earth’s life [Cleland and Chyba, 2002; SchulzeMakuch and Darling, 2010], nevertheless there is no available data beyond that of our own ecosystem. Therefore it is most appropriate to evoke the fundamental requirements of life, such as liquid water, to search for examples beyond the Earth. Not all liquid water is habitable however, and identifying inhospitable liquid water can constrain the search for potentially habitable environments on other planets. A pressure-temperature model of liquid water and life can be used to refine and focus the searches for life on Earth and elsewhere. In this chapter the P-T model of the Earth and its water chemistries developed in Chapter 3 is utilised, and compared to the range of P-T conditions that occur within the biosphere.
This
allows the question of whether there are regions of the Earth where liquid water is uninhabited to be considered, by identifying where the P-T space of liquid water in the Earth is not overlapped by the P-T space of known life (Figure 4-1). Such a region would imply that life is restricted from some liquid water environments by conditions of temperature, water activity, pressure, nutrients or energy beyond the tolerance limits of all terrestrial organisms. In turn, this would imply that the strategy of follow the water is not sufficient in the search for life on other planets and can be refined, with physical and chemical modelling playing a significant role.
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Figure 4-1: Liquid water, Earth and life. In Chapter 3 the extent of liquid water on Earth was modelled, by studying the degree of overlap in pressure-temperature space. In this chapter the phase space of life will be modelled, and utilised to consider the extent of the biosphere on Earth and the fraction of inhabited liquid water environments.
4.1.
Development of the biosphere model
The phase model for the biosphere was constructed through identifying the extreme pressures and temperatures (P-T) of environments known to be inhabited by life, listed in Table 4-1 and Table 4-2. The maximum values of P and T define the known life-space of Earth organisms. The term ‘life box’ has been used to indicate the overall range of conditions that an organism experiences [Wharton, 2002].
Here, the term ‘life space’ refers to the full range of P-T conditions
experienced by life on Earth. In the process of quantifying this region, the true P-T space of life can be underestimated due to the effects of sampling bias. Undersampling of the life-space may occur both through incomplete surveying of the published literature (it is not practical to compile every measured pressure and temperature value from inhabited environments), and through biased sampling of the biosphere (not every inhabited environment has been studied, for example, deep environments within the crust).
78
Phase Diagram of Earth’s Biosphere
The ‘maximal’ life-space is defined as encompassing all P-T values within the boundary of extreme P-T conditions experienced byorganisms on Earth (Figure 4-2). It extends from the hottest high pressure inhabited environment, to the coldest low pressure inhabited environment, and includes all P-T values within that range. The maximal life space is bigger than the known life space as it also includes environments that are within the range of P-T conditions known to be hospitable to some organisms on Earth, but for which a specific example was not found in the literature.
This is done to compensate for the aforementioned
potential undersampling of the known life space.
However, the available
environments for life are limited by the P-T space of the Earth, so an estimate of the ‘actual’ life space for terrestrial organisms is the ‘maximal’ life space mapped onto the pressure-temperature conditions within the Earth.
Figure 4-2: Schematic of the method for constructing the biosphere phase model. Blue dots represent pressure-temperature data obtained from the literature. Green represents the modelled P-T space of life. Brown represents the P-T space of the Earth derived in Chapter 3. The darker green shading indicates regions of Earth that were added to the known life space to produce the maximal life space. These are environments that plausibly have life but have not been sampled (4.1.1). The best estimate of the actual/true life space of active and dormant life on Earth is derived by taking the maximal life space and mapping it onto the phase space of Earth.
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4.1.1. Extrapolating to define maximal life space In modelling the inhabited regions of P-T space, it was assumed that the upper temperature limit to life of 122 ˚C [Takai et al., 2008] and the low temperature limit for active life in liquid water of -20 ˚C [Junge et al., 2004] are both valid over a broader range of pressures than has been fully explored by microbiologists. This is a valid assumption, as a strong link has not yet been established between the extreme temperature and extreme pressure adaptations of organisms. In other words, an organism that grows at 122 ˚C at ~ 107 Pa (point 4, Table 4-1) may be just as adapted to grow at the same temperature at 108 Pa (point 5). The link between barophilic and thermophilic adaptation is still unclear [Rothschild and Mancinelli, 2001], although many hyperthermophiles are also barophiles [Marteinsson et al., 1997, 1999; Takai et al., 2000] (2.6.1). This correlation may however be due to the elevation of the boiling point temperature of liquid water under high pressure (2.4.2) and hence the availability of higher temperatures for life in deep ocean or deep crustal environments. It is also likely that many of the adaptations that allow organisms to thrive in hot environments, such as proteins able to maintain their structure and stability at high T, may also enable an elevated pressure tolerance or vice-versa [Holden and Baross, 1995; Pledger et al., 1994; Robb and Clark, 1999] (2.7.1). Organisms that are active in low pressure high altitude environments on Earth also show adaptation to desiccation stress [Yang et al., 2009a, 2009b], similar to that encountered by organisms in sub-zero liquid water environments [Steven et al., 2006]. Low pressure however has not been observed to fundamentally limit activity and growth (2.6.2), and hence an organism that grows at -20 ˚C and 107 Pa (point 8, Table 4-1) is likely to be just as adapted for growth in a liquid water environment at -20 ˚C and ~ 104 Pa (point 1, Table 4-1). Table 4-1 indicates which P-T environments for active life have been extrapolated to higher or lower pressures.
Based on the discussion above,
extrapolated environments were determined by: (i) predicting that life adapted to extremely low temperatures can maintain that adaptation at a variety of lower pressures; and (ii) that life adapted to extreme high temperatures can maintain their adaptation at higher pressures.
Although plausible based on the links
between thermophlic/barophillic and psychrophilic/osmpophilic adaptation (see also 2.6 and 2.7), these biosphere environments should be targeted for biological sampling to determine their habitability and verify the predictions of this model.
80
Phase Diagram of Earth’s Biosphere
4.1.2. Active & dormant life Environmental stresses such as extreme temperature, water desiccation and starvation, can cause some organisms to enter a state of extremely low metabolic activity known as dormancy. This state, including metabolism as low as 5 % of normal [Wharton, 2002], is reversible. Dormant organisms are able to resuscitate after environmental change re-introduces tolerable conditions for growth and reproduction [Lennon and Jones, 2011], even after extremely long periods of dormancy [Nicholson et al., 2000]. In this work, the distinction is made between PT conditions that support active life, and those where only dormant life has been identified. As organisms can neither grow nor multiply in a dormant state and hence cannot complete their life cycle, it is difficult to consider such environments as being ‘habitable’ if no active life is present. However, it is plausible that within the phase space of dormant life there are organisms which can remain active but have not yet been sampled, hence for completeness dormant life environments are included within the estimated life space. Furthermore, environments which are known to induce dormancy could provide important targets in the search for living or preserved microorganisms on extraterrestrial bodies [Abyzov et al., 2006]. In Table 4-1 and Table 4-2 the pressure-temperature values that have been used to derive the boundary region of active and dormant environments, respectively, are given (plotted below). They include both the environments where life has been found and the extrapolated environments where life is expected, given that there is currently no conclusive evidence that high or low temperature adaptation is correlated with pressure. For interest, Table 4-3 provides some further extreme pressure habitats – at high altitude and deep subsurface – that are also included within the life region. environments.
These habitats help to define the range of habitable
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Table 4-1: Extreme active life environments Key*
Description
1
Mt Everest summit
2
Thermal spring/vent (Pork Chop Geyser, Yellowstone National Park, USA) Shallow hydrothermal vents
7.7×104 -2.29
97
2×105 0.01
122
Ocean hydrothermal vent where highest temperature life found (Finn Vent, Mothra field, Pacific Ocean) Marianas trench, deepest part of ocean (Challenger Deep, Pacific Ocean)
2.1×107 2.27
122
1.32×108 10.92
122
Deepest active life in crust from borehole (Gravberg, Sweden) Marianas trench (Challenger Deep, Pacific Ocean) Arctic ice core, lowest temperature active life.
1.55×108 5.28
75
1.32×108 10.90
2
[Kato et al., 1998; Takami et al., 1997; Todo et al., 2005; Yayanos et al., 1981]
107 2.75
-20
[Junge et al., 2004]
3
4
5
6
7 8
Pressure (Pa) Temp. Depth (km) (˚C) 3.3×104 -20 -8.85
*Green shading indicates extrapolated conditions
Reference Life has not been fully searched for at this location. This location is extrapolated from point 8 to lower pressures. [Guidry and Chafetz, 2003]
Thermophillic life has not been fully searched for at this location. Life has not yet been found in environments hotter than 100 ˚C at this depth [Amend et al., 2003; Maria ProlLedesma, 2003; Stetter, 2006], despite the existence of environments > 122 °C [Maria Prol-Ledesma, 2003] . This point is extrapolated from point 4 to lower pressures. [Kashefi and Lovley, 2003; Schrenk et al., 2003; Takai et al., 2008]
Thermophillic life has not been fully searched for at this location. Life has not yet been found in environments hotter than ~3 ˚C at this depth [Todo et al., 2005], despite the existence of environments > 122 °C. This point is extrapolated from point 4 to higher pressures. [Szewzyk et al., 1994]
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Phase Diagram of Earth’s Biosphere
Table 4-2: Dormant life environments Key Description
Pressure (Pa)* Depth (km) 107 2.75
Temp. Reference (˚C) -40 [Abyzov et al., 1998; Price and Sowers, 2004]
9
Vostok ice core life
10
Greenland open mine life
105
-30
11
Mt Everest summit
3.3×104 -8.85
-40
12
Bacterial spores (atmosphere) Non-spore forming eubacteria (atmosphere) Bacterial spores (atmosphere)
5.53×103 -20
-20
2.5×102 -41
-56
2 -77
-69
13
14
Metabolic activity
Cell activity after incubation; interpreted as cells in situ were able to maintain vital functions but not reproduce. [Langdahl and Sampled at warmer Ingvorsen, 1997] temperatures so metabolic activity at low temperature limit is unknown. Life has not been fully searched for at this location. This point is extrapolated from point 9 to lower pressures. [Wainwright et al., Dormant (spore) in 2003] atmosphere but cultured in laboratory. [Griffin, 2008] Growth after extended incubation. [Imshenetsky et al., 1978]
Dormant (spore) in atmosphere.
Table 4-3: Further extreme habitats for active life Description
Temp. (˚C) 100
Reference
76
[Onstott et al., 1998]
-6
[Alekhina et al., 2007]
Cloud life in water droplets
Pressure (Pa) Depth (km) 8.1×107 1.6 subocean; 6.1 subsurface 3.8×107 2.8 3.2×107 3.6 6.9×104 -3.1
-5
Highest altitude active life
5×104 -5.6
-20
[Sattler et al., 2001; United States Committee on Extension to the Standard Atmosphere, 1976] [Gangwar et al., 2009]
Deepest sub-ocean life Deep life in sedimentary rock Deep life in ice
[Hamilton, 1976; Roussel et al., 2008]
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4.1.2.a.
Volume calculations
To calculate volume fractions of the Earth in 4.2, such as the fraction of the Earth’s volume that is known to be inhabited by life, the following equation was used: ሾ ሿିሾ௭ሿ ቁ ሾ ሿ
Ψܸ ݄ݐݎܽܧ݈ൌ ቀ
Equation 4-1
ସ
where ሾሿ ൌ ଷ ߨሺݎ െ ݖሻଷ is the volume of the Earth not included in a spherical
shell of thickness z extending from the surface, z = depth (km) and ݎ ൌ ͵ͺ km is the radius of the Earth. To calculate volume fractions of the hydrosphere, such as the fraction of liquid water on Earth that is inhabited by life, the following equation was used: ሾ ሿିሾ௭ሿ
Ψܸ ݎ݁ݐܹ݈ܽൌ ቀሾ ሿିሾ௭ ሿቁ
Equation 4-2
ೢ
ସ
where ሾ୵ ሿ ൌ ଷ ߨሺݎ െ ݖ௪ ሻଷ is the volume of the Earth not included in the spherical shell of thickness ݖ௪ ൌ Ͳ km that can support liquid water from
pressure and temperature constraints (see 3.2.2). zw is the maximum extent of the hydrosphere. The relationship between depth and pressure values in Table 4-1, Table 4-2 and Table 4-3 is given in Equation 3-1. Densities are predominantly provided in Table 3-4. Additional densities include ocean water (1025 kg/m3), ocean floor sediment (2300 kg/m3) and water ice (917 kg/m3).
84
Phase Diagram of Earth’s Biosphere
Figure 4-3: Estimated phase space of Earth life superimposed on the P-T diagrams of Earth water from Figure 3-3.
Numbers 1-8 circumscribe the estimated region
occupied by active life. Numbers 9-14 circumscribe the estimated region occupied by dormant life.
The maximal P-T space of life has been mapped to Earth’s
environments (Figure 3-5) in this plot. For example, the width of the dormant region is constrained not only by Table 4-2 but also by the P-T environments of Earth’s atmosphere (see Figure 4-4).
4.2.
Results: the pressure temperature representation of the
biosphere The phase space of the biosphere is shown in Figure 4-3, superimposed on the phase diagram of Earth’s water chemistries - ocean water and the maximum liquid range of brines and thin films (3.1.1). All known examples of life on Earth inhabit the pressures and temperatures shaded in green, with the cross-hatching indicating environments where only dormant life has been found. In Figure 4-4 the phase space of the biosphere is superimposed on the phase space of the Earth. Below the fraction of liquid water environments on the Earth that may be uninhabited is quantified.
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Table 4-4: Relevant temperature isotherms and their depths from Earth P-T model Description
Temp. Mean pressure
Mean depth
% Volume
% Volume
(˚C)
(km)* [range]
Earth**
Earth within
(Pa) [range]
hydrosphere **# Current max.
122
temp. for life Proposed max.
150
temp. for life Proposed max.
250
temp. for life
1.1×108
4.0
[105 – 4.0×108]
[0- 14.6]
1.8×108
6.6
[105 – 6.3×108]
[0-23.0]
2.8×108
10.2
[105 – 1.3×109]
[0- 46.4]
0.2 (0.6)
5.8 (16.6)
0.3 (0.7)
9.5 (19.9)
0.5 (0.9)
14.7 (25.5)
*See Equation 3-1 and Table 3-5 for details of calculation **Assuming spherical symmetry and using Equation 4-1. In the curved brackets, 9 km is added to the Earth’s radius to include the estimated maximum altitude of life above the surface (Table 4-2). #Assuming spherical symmetry and that the hydrosphere extends to a maximum depth of 70 km (3.2.2), and using Equation 4-2. The volume therefore is the volume of the Earth with liquid water - the 70 km thick shell extending from the surface.
On the next page: Figure 4-4: Superposition of inhabited environments from Figure 4-3 on the phase space of the Earth and water from Figure 3-5. Based on current knowledge, life does not inhabit the full range of Earth’s environments where liquid water is available. The highest temperature environment of active life, currently 122 ˚C, excludes life from the hottest and deepest water. The orange and yellow lines represent the range of depths that life could inhabit if its high temperature threshold were found to be 150 ˚C and 250 ˚C respectively (2.7.1). The red asterisk represents the prediction of the hottest and deepest water on Earth (3.2.2).
Figure 4-4
86
Phase Diagram of Earth’s Biosphere
87
4.2.1. Extent of the biosphere The known biosphere encompasses the full extent of the ocean and the continental crust through to at least one location of ~ 5.5 km depth. Both the maximum depth of life in the oceans (11 km at the Marianas Trench [Kato et al., 1998]) and within the crust [Szewzyk et al., 1994] occur at a similar pressure (points 6 and 7 in Table 4-1). These locations reflect the maximum extent of the known biosphere. The environments are limited by sampling bias however, as life has been found in the crust, ocean, and sub-ocean sediments as deep as microbiologists have searched for it. Hence, given that there are extensive regions within the Earth at higher pressures than life has been found, but within the temperature thresholds of life (Figure 4-4), this implies that the apparent high pressure limit of life is a selection effect. Furthermore, this conclusion is supported by experimental studies in which barophiles were subjected to the pressures encountered down to ~ 50 km depth within the crust and yet remained alive and viable [Margosch et al., 2006; Sharma et al., 2002] (2.6.1), implying that high pressure itself is not a limiting factor for biology. Assuming the availability of liquid water and nutrients, it is more likely that high temperature and pore space are the limiting factors for the deep biosphere [Gold, 1992]. The maximum plausible extent of life within the Earth can be quantified by the depth of the 122 ˚C isotherm. Currently the hottest environment supporting life is 122 ˚C [Takai et al., 2008]. As there is no evidence for a pressure limit to life it is expected, as microbiologists explore deeper within the subsurface, that the depth of the biosphere is determined by the depth of this isotherm. From the pressuretemperature model of the Earth (Figure 3-4) liquid water is not expected to become a limiting factor for life until much greater depths (between 60-70 km depending on pore space availability; 3.2.2).
The intersection of the 122 ˚C
isotherm with the mean geotherm of the Earth model in Figure 4-4 occurs at a depth of 4 km (Table 4-4), and is within the liquid phase space of ocean water. There is significant variation in the depth of the isotherm however, as evidenced by active life in only 75˚C conditions at a greater depth of 5.5 km [Szewzyk et al., 1994]. From Figure 4-4, the depth of 122 ˚C varies between 0-14.6 km, with the variation due primarily to heterogeneity in the crustal heat flux. The smaller
88
Phase Diagram of Earth’s Biosphere
bound occurs at the surface in environments such as hot springs. The larger bound, which could occur in regions of thick continental crust with low radiogenic abundance and heat flow, will be testable through future biological drilling missions [Boutt et al., 2007; Moser et al., 2010]. As liquid water has been sampled in the crust at a depth of 11 km [Jodicke, 1992], water at ~ 15 km is highly plausible. The current maximum temperature of life may be the result of the limited exploration of high temperature environments in hydrothermal vents.
It is
plausible that in the future the real temperature limit of life will be found to be higher- perhaps approaching 150 ˚C [Segerer et al., 1993] or even 250 ˚C [Baross and Deming, 1983; White, 1984] (2.7.1). These two scenarios, summarised in Table 4-4, would result in the mean and range in depths for the lower limit of the ଷǤଶ biosphere extending to 6.6േ଼Ǥ Ǥ km (orange line in Figure 4-4) and 10.2േଵǤଶ km
(yellow line) respectively. Ocean liquid water is stable over both these depth ranges and so neither of these temperature limits can be discredited.
Both,
however, would result in a significant extension of the depth of the biosphere and the phase space of life.
Figure 4-5: Deep hot life in restricted pore space. Bacteria isolated from 2.7 km deep gold mine at in-situ temperatures of 75 ˚C and pressures of ~ 3.2×107 Pa. It is anaerobic, thermophillic, halo-tolerant, and surviving in subsurface rocks with only 0.04 vol% porosity [Onstott et al., 1997, 1998].
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4.2.1.a.
Pore space & permeability
An additional limiting constraint on the base of the biosphere is restricted access to liquid water, measured by pore space and, closely related, permeability. Despite subsurface P-T conditions that allow water to be liquid within the temperature thresholds of Earth life, the existence of both liquid water and life will depend on the availability of pore space. Pore space of surface materials ranges typically from 10-60% [MacCary and Lambert, 1962]. From the discussion in section 3.2.2, pore space in crustal rocks is expected to be extremely limited to much less than 1 % above pressures of 350 MPa [Abdulagatov et al., 2006] and absent in any rocks that have experienced temperatures over 350 ˚C [Hyndman and Shearer, 1989]. This equates to extremely limited pore space between ~ 12 km (from the pressure constraint and the mean Earth model geotherm) and ~ 63 km depth (from the temperature constraint, Table 3-6) and hence may effectively limit the lower boundary of the biosphere to being somewhere within this depth range. Life may have a minimum pore space requirement, although active life has been isolated from rock with extremely low pore space as in Figure 4-5. It is likely however that the connectedness of pore space, which governs the flow of fluid, is also fundamental.
For example, permeability is an important control on the
distribution of microbial communities in sea ice [Thomas and Dieckmann, 2002]. Permeability is related to porosity in a complex way that is not well known [Rezaee et al., 2006]. Permeability decreases as a power law with increasing depth [David et al., 1994] and also varies widely with the type of rock, mineralogy, fracturing and grain cementation [Ingebritsen and Manning, 1999; Wang et al., 2009]. Permeability is significantly restricted below the brittle-ductile transition at ~ 1015 km depth [Manning and Ingebritsen, 1999], due to the decreased fracturing and porosity of the rocks in situ. Furthermore, borehole data indicates that oceanic crust may become impermeable at much smaller depths of < 1 km [Anderson et al., 1985]. Considering the above discussion, pore space and permeability place a similar threshold on the deep boundary of the biosphere, with restricted access to fluid occurring between ~ 10-15 km but varying strongly with the substrate and potentially allowing pore fluid as deep as 63 km [Hyndman and Shearer, 1989]. For
90
Phase Diagram of Earth’s Biosphere
example, despite being nearly at the depth of the brittle-ductile transition, fluidfilled fractures have been identified through drilling in rocks at 9.1 km depth [Huenges et al., 1997]. The mean depth of both the 122 ˚C and 150 ˚C isotherms fit comfortably within the depth constraint of pore space and permeability, and hence the temperature limit for life may render the latter thresholds irrelevant.
4.2.2. The subzero biosphere Figure 4-3 depicts active and dormant life in cold environments outside of the liquid temperature range of Earth’s oceans.
This life exists in briny liquid
inclusions and thin films within permafrost and ice, such as in Figure 4-6. Microorganisms have been found growing and metabolising in subzero environments within pockets of liquid water [Priscu and Christner, 2004] provided that the temperatures remain above -20 ˚C [Junge et al., 2006] (2.7.2) and the activity of the liquid is above 0.6 [Grant, 2004] (2.5.3). This subzero water is maintained as a liquid due to the presence of salts which lower the freezing point, in exchange for lowering the liquid activity. From the discussion in section 2.5, it is likely that the apparent low temperature limit observed for active life is due to the water activity limit and the absence of liquid water environments at lower temperatures on Earth with a high enough activity for life. The importance of the water activity threshold can be tested in laboratory studies, as certain brines (which are not common in Earth’s polar environments) maintain a high activity at their freezing point of -30 ˚C. From Figure 4-4, a significant region of the P-T space of the Earth in the crust and atmosphere occurs at temperatures below -20 ˚C but within the maximum liquid range. Earth’s permafrost contains typically < 7 wt% unfrozen water [Gilichinsky et al., 1993], with ~ 2 wt% typical of -20 ˚C [Rivkina et al., 2000] but reaching as high as 15 wt% within some clay soils [Dillon and Andersland, 1966]. Similar unfrozen water contents occur within arctic and glacial ice [Price, 2000]. Hence if life is limited by water activity to environments warmer than -20 ˚C, this indicates a significant fraction of uninhabited liquid water environments in the Earth’s crust. The distinction between a low temperature threshold and a low water activity threshold for life is significant for quantifying the habitability of other planets. Different concentrations of salts in different planetary environments could result in the low activity threshold of 0.6 for life occurring at a much lower or much higher temperature than -20 ˚C.
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Figure 4-6: Deep, icy life from Antarctic ice core4. Cell activity could be measured after incubation, indicating that cells in situ were dormant and able to maintain vital functions but not reproduce. In situ temperatures were -40 ˚C at 3.5 km depth [Abyzov et al., 2001; Price and Sowers, 2004].
4.2.3. The ambiguity of cold atmospheric life From Table 4-3 there is no confirmed active life at pressures less than 5×104 Pa (5.6 km altitude), however given that the conditions fall within the temperature thresholds of life there is potential for active life at 30 kPa (~ 9 km altitude). The discussion in 2.6.2 indicated that organisms are not fundamentally low pressure limited [Kerney and Schuerger, 2011]. Rather the challenge for high altitude life is the low availability of liquid water and hence the extreme stress of desiccation [Sattler et al., 2001]. A further challenge faced by atmospheric life and its capacity to remain actively metabolising is the transient nature of liquid water environments due to the continual movement of air currents. For example, at the triple point of ocean water in Figure 4-4 at ~ 25 km altitude, the conditions are likely dry (even though liquid water is stable) due to the removal of liquid water from the air as it migrates above the troposphere [Peter et al., 2006]. Droplets of liquid water within clouds can contain nutrients [Womack et al., 2010] and persist throughout the troposphere which has an average thickness of ~ 15-20 km [Randel et al., 2004].
4
Although this range of altitudes provides an estimate for the
http://science.msfc.nasa.gov/newhome/headlines/ast12mar98_1.htm
92
Phase Diagram of Earth’s Biosphere
maximum possible extent of liquid water above the surface of Earth, the conditions are most likely too cold for active life. Table 4-5: Volume estimates for the known biosphere and plausible hydrosphere Description
Depth* (km)
Fraction of Earth’s volume* (%)
Earth Maximum depth extent of the hydrosphere Maximum depth extent of the hydrosphere considering pore space restrictions Maximum depth of life in crust Maximum depth of life in combined ocean and oceanic crust** Mean depth of 122˚C isotherm Mean depth of 150˚C isotherm Mean depth of 250˚C isotherm
6378 (6387) 70 (79)
100 3.3 (3.7)
Fraction of Earth’s volume within hydrosphere* (%) ---100 (100)
63.4 (72.4)
3.0 (3.4)
100 (100)
5.5 (14.5) 6.1 (15.1)
0.3 (0.7) 0.3 (0.7)
7.9 (18.5) 8.8 (19.3)
4 (13) 6.6 (15.6) 10.2 (19.2)
0.2 (0.6) 0.3 (0.7) 0.5 (0.9)
5.8 (16.6) 9.5 (19.9) 14.7 (24.5)
*Assuming
spherical symmetry and using Equation 4-1. Brackets indicate the addition of 9 km to the Earth’s radius to include the estimated maximum altitude of life above the surface (Table 4-2). See Equation 4-1 and Equation 4-2 for details of calculation. ** Depth of the Marianas trench (11 km) is not included as it is an outlier and would skew volume calculations. The mean ocean depth is 3.8 km and the majority of ocean floor is shallower than 6 km.
4.2.4. Uninhabited liquid water & its implications for the search for life The superposition of the pressure-temperature space of the biosphere onto all environments within the Earth allows the volume of the Earth that is inhabited by life to be quantified. Here only active life is considered as being indicative of habitable environments. In the P-T model the biosphere extends from ~ 9 km above the mean surface of the Earth to at least one location ~ 5.5 km deep within continental crust. Life has also been found to 6.1 km depth through ocean and subocean sediments [Roussel et al., 2008] and occurs through the full extent of the oceans [Kato et al., 1998]. Hence for the purposes of this calculation the mean depth of life through the oceans and oceanic crust can be estimated as ~ 6.1 km. Using Equation 4-1 and assuming oceans occupy 70 % of the surface and continents occupy 30 %, the known biosphere occupies ~ 0.7 % of the volume of Earth with 9 km of atmosphere (Table 4-5). This value is an upper limit, as not all environments in the troposphere are inhabited [Amato et al., 2005] and given the
93
average continental and oceanic crust heat flow (Table 3-4) the 122 ˚C isotherm would generally be encountered shallower than 5-6 km depth. For example, life was absent from a 4.1 km borehole where in situ temperatures exceeded 118 ˚C [Huber et al., 1994]. In summary, after 4 billion years of evolution, the Earth’s biosphere has extended into < 0.7 % of the volume of the Earth, leaving > 99.3 % uninhabited, given the current exploration of subterranean life. The inhabited volume of the Earth could extend to as high as 0.9 % if active life is found in environments up to temperatures of 250 ˚C (Table 4-5). The fraction of liquid water environments that are inhabited by life can also be quantified. Without data on the variable volume of liquid water within the Earth with depth, the volume of Earth where liquid water is stable can be used, as this is the best estimate of where liquid water exists (3.2.2).
The computation is
analogous to that above except instead of using the volume of the entire Earth the volume of the Earth with liquid water, the hydrosphere, is used (Equation 4-2). As the pressure and temperature environment of the deep Earth allows liquid water to be stable to a depth of ~ 70 km (red asterisk in Figure 4-4), the volume used is a spherical shell of this thickness.
Again a mean depth of 5.5 km for life in
continental crust; 6.1 km for life in combined ocean and oceanic crust; and a thickness of 9 km for the biosphere in the atmosphere is used.
Therefore from
Table 4-5, the known extent of the biosphere occupies ~ 19 % of the volume of the Earth within the hydrosphere. Hence 81 % of the volume of Earth where liquid water is stable and most likely exists, is not known to harbour life. Even if life is found active in environments up to 250 ˚C, this estimate will only increase to 24.5 % of the volume of Earth with liquid water being hospitable to life, based on the mean depth of the 250 ˚C isotherm (Table 4-5). The distinctions made amongst the biosphere, hydrosphere and the inhabited volume of the Earth can be visualized in Figure 4-7.
Although the volumes
discussed above and in Table 4-5 are estimates, they provide a first quantification of the extent that life has expanded throughout the Earth (new results soon to be published will make an interesting comparison [Méndez, 2012]). Furthermore, the results quantify the extent to which life has been limited from the deep liquid water environments of the Earth by high temperature. Although all known life on
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Phase Diagram of Earth’s Biosphere
Earth requires liquid water, life is excluded from a large percentage of the phase space of liquid water given the current understanding of the high temperature limit of life. It is expected that the volume estimates of the biosphere, and the fraction of Earth within the hydrosphere that is inhabited by life, will alter if life is found significantly deeper within the crust (and when the minimum pore space and permeability for life becomes constrained). More relevant to the search for life on other planets (such as Mars, Europa, Titan, etc.), is the restriction of life from cold liquid water environments. Although this is difficult to quantify as a volume of Earth, instead the fraction of the pressure-temperature space of briny liquid water and thin films that is occupied by life, shown in Figure 4-3, can be considered. The fraction of the area of liquid water phase space [log P×log T] that is occupied by life is 36 %: 30 % is occupied by active life and the remaining 6 % by dormant life. As the majority of the P-T space of this liquid water overlaps with environments on the Earth (Figure 4-4), the 64 % of the phase space of Earth liquid water not known to be occupied by life represents a significant constraint on the biosphere.
4.2.5. Estimating the amount of uninhabited water Given current knowledge on environments that harbour life, uninhabited liquid water environments on Earth occur predominantly at temperatures above ~ 122 ˚C and depths below ~ 4 km in continental crust. In previous sections, the biosphere was quantified as a fraction of the volume of Earth within the hydrosphere- the 70 km thick shell from the surface of Earth where the P-T conditions allow for liquid water. Figure 4-8 clarifies the distinction between these points. A substantially different approach can also be taken however, to quantify the amount of liquid water on Earth that is potentially inhabited in terms of the known water budget of the planet.
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Figure 4-7: Schematic of Earth's hydrosphere and biosphere. In the context of the volume of Earth, the regions shown in Figure 4-4 are plotted: water and life (outer shell), water and no life (middle shell), and no water and no life (inner shell). The thickness of the known biosphere is ~ 6 km in the crust and ~ 9 km in the atmosphere (15 km total), which represents 0.9 % of the volume of the Earth. Liquid water exists to a depth of ~ 70 km in the lithosphere and through 9 km of atmosphere (79 km total), which represents 3.7 % of the Earth’s volume. 96.3 % of the volume of the Earth has no liquid water and hence is uninhabitable by life. From 3.2.3 there is ~ 1.36×109 km3 of liquid water on the Earth within the oceans and as groundwater within the crust. Oceans comprise ~ 98.5 % of this volume, with the remaining ~ 2×107 km3 existing as groundwater km3 [Foster and Chilton, 2003; Lehr and Lehr, 2000; Rogers and Feiss, 1998]. The oceans are known to be inhabited by life throughout the full extent of their explored depths [Kato et al., 1998], and so only the habitability of the groundwater need be considered. From 3.2.3 again, ~ 9.5×106 km3 of groundwater is shallower than the average 4 km depth of the 122 ˚C isotherm. Therefore by subtracting from the total amount of liquid water on Earth indicates that an estimated 1.05×107 km3 of groundwater may be inhospitable to life, reflecting 0.8 % of the Earth’s total liquid water budget.
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Phase Diagram of Earth’s Biosphere
Hence, the vast majority Earth’s water budget: ~ 99.2 % of liquid water and ~ 47 % of groundwater, is within the temperature thresholds of life (Figure 4-8).
Figure 4-8: Quantifying uninhabited water. Considering the environments on Earth that have liquid water, two ways of visualising what fraction of these environments are inhabited by life are plotted. Left: Of the 3.5 % of the volume of Earth where liquid water exists, 19 % is inhabited, 81 % is uninhabited. If the upper temperature limit for life is found to be 250 ˚C (dashed line) then the inhabited fraction increases from 19 % to 25 %. Right: Of the amount of liquid water on Earth, 99 % is inhabited. At least 1 % of Earth’s liquid water budget is uninhabited.
4.3.
Summary
Based on the currently known extent of life, the Earth’s biosphere inhabits only a small subset of the full ~ 70 km depth extent of liquid water within the Earth. The current high temperature limit for life is 122 ˚C, however this limit may not be fundamental for all life as there is potentially a substantial amount of liquid water (< 53 % of groundwater budget) at higher temperatures that has not been fully searched for life. If 122 ˚C is a real limit, it reflects a severe restriction of the biosphere, with only ~ 19 % of the volume of Earth with liquid water supporting life. Additionally life may be restricted from cold near surface regions in ice and permafrost by low water activity, despite liquid water being available (brines and thin films) at 3.5-15 vol% in environments colder than the minimum known
97
temperature supporting active life of -20 ˚C. All examples of life found at pressures less than ~ 10-4 Pa have been classified as dormant indicating that life at high altitude may be limited by severe desiccation and nutrient stress. In Chapter 6 the P-T model of the Earth’s biosphere and the limitations on life within liquid water will be applied to the modelled environments in the subsurface of Mars.
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Phase Diagram of Earth’s Biosphere
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5. LITERATURE REVIEW & BACKGROUND: LIQUID WATER ON MARS
‘In absences you grew more beautiful, more poisonous. You were an attar of orchids in the stemming night, where passion, like a shark having found a bloodstream, murders other senses. Only taste preserving, buckling into itself, finding the blood its own, a small wound first. But as the shark unravels the belly tatters in the long throat’s tunnel. And knowing this, the night still seems a richness, a gauntlet of desires ending in peace, I would still be part of these allurements, and to my arms I would take in the darkness, blessed and renamed by pleasure. But the light, The light, And so into the shadow, and not your shadow, but the eager grayness expecting light, I ride the storm away.’ - Margaret Weiss and Tracey Hickman, ‘The Dragonlance Chronicles (Dragons of Spring Dawning)’
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Literature Review & Background: Liquid Water on Mars
Mars is known to have extensive water ice on its surface and in its shallow subsurface at mid- to high latitudes and a global water vapour abundance, and significant erosion by liquid water in its past [Head et al., 1999]. There is also strong evidence for past long-term liquid water environments in the form of Noachian-aged phyllosilicate and carbonate deposits; and Hesperian- and Amazonian-aged hydrated silicas and sulphates, formed from standing water interacting with basaltic crustal rocks [Ehlmann et al., 2008, 2009; Gendrin et al., 2005; Hamilton et al., 2008; Murchie et al., 2009; Mustard et al., 2008; Poulet et al., 2007; Wray et al., 2009]. Mars has experienced dramatic climatic changes [Laskar, 2004], with its varying obliquity resulting in the present day thin atmosphere and the loss and redistribution, through freezing and evaporation, of past reservoirs of liquid water from the surface [Clifford, 1993]. Despite this, Mars is still estimated to have a significant water inventory of at least 106 km3 of water ice [Christensen, 2006], and unknown volumes of liquid water. The presence of water on Mars and the likelihood of liquid water, identifies Mars as having strong potential for hospitable environments able to support life. In this review, the previous work on identifying current liquid water environments on Mars will be discussed, and an assessment of the morphological evidence for water and its stability under the current martian climate will be provided.
5.1.
Water stability
Although extensive water ice environments have been identified on Mars, and soil temperatures in the low to mid- latitudes are frequently above melting [Smith et al., 2006], the difficulty locating water in the liquid phase is the transience of any water exposed to the martian atmosphere. Due to the low partial pressure of water vapour, most surface water environments on Mars exist for only seasonal or diurnal timeframes. The duration of martian water is discussed below.
5.1.1. Ice Across most of the martian surface the total atmospheric pressure (Ptot) is less than the triple point of water [Haberle et al., 2001] and hence ice and vapour are the dominant phases of water. The occurrence of water ice can be predicted from the temperature and partial pressure (Ppart) of water in the atmosphere. On any
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planet, the temperature at which water ice will sublimate or water vapour will condense is given by the sublimation curve in the phase diagram of water, at the corresponding Ppart of water [Silberberg, 2003]. In the martian atmosphere, Ppart varies between ~0.01–2 Pa [Melchiorri et al., 2007, 2009; Smith et al., 2009b] but is typically ~0.1 Pa [Jakosky and Phillips, 2001; Whiteway et al., 2009], shown by the red shading in Figure 5-1. For comparison, a typical Ppart at sea level on Earth is much higher at 300 Pa [Goody and Belton, 1967]. From the phase diagram of water at the mean Ppart of the martian surface (0.1 Pa) the transition from solid to vapour occurs at point A, at 200 K (-73 ˚C, Figure 5-1). This is referred to as the martian frost point. From the variation in Ppart at the martian surface the frost point temperature varies between 190-210 K (-83 ˚C to -63 ˚C). At temperatures above the frost point any ice exposed to the atmosphere will not be stable and will sublimate [Mellon and Phillips, 2001]. At temperatures below the frost point, water vapour will condense onto the surface and diffuse into the soil to condense as ice [Mellon and Jakosky, 1993]. In general, any amount of water vapour exceeding the equilibrium vapour pressure (Psat) at a given temperature is converted into ice [Jakosky, 1983; Richardson and Willson, 2002], which in turn decreases the vapour pressure Ppart. Condensation will cease when Ppart reaches the sublimation curve (i.e. when Ppart = Psat). Saturation is observed on Mars when the temperature decreases at a given Ppart [Smith et al., 2009b]. For example, point A in Figure 5-1 indicates saturated atmosphere at 200 K (-73 ˚C) but an unsaturated atmosphere at 250 K (-23 ˚C). The condensation of water vapour into ice on the surface is not instantaneous however, so the atmosphere can remain supersaturated for a period of time, as has been observed in the martian northern lowlands [Maltagliati et al., 2011]. Frost can form directly from vapour without a transient liquid phase [Schorghofer, 2005].
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Literature Review & Background: Liquid Water on Mars
Figure 5-1: Partial pressures of water on Mars. The red crosshatching indicates water vapour conditions in Mars’ atmosphere. Point A indicates the frost point at a typical Ppart (200 K, 0.1 Pa).
The arrows indicate the effect of reducing the
atmospheric T without changing the Ppart. For example, if the temperature is increased from point A the atmosphere becomes unsaturated, and ice will be vapourised until Ppart increases to reach the vapour curve at the increased temperature (Psat). Point B indicates the boiling condition on Mars: Ppart = Ptot. As the partial pressures do not reach point B, boiling cannot occur on Mars. Point C indicates a typical indoors environment on Earth. For reference, the green and red circles indicate a typical Ptot, T condition at sea level/zero elevation on Earth/Mars respectively.
This has implications for the duration of water exposed to the
atmosphere. The stability and depth of ice in the subsurface depends not only on the atmospheric Ppart, but also on the temporal variability of the soil temperaturespecifically, the frequency and depth of temperatures above the frost point. This in turn depends strongly on the response of the materials to solar and geothermal heating.
Considering near surface materials in contact with the martian
atmosphere,
the
subsurface
temperature
environment
is
determined
predominantly by the thermal inertia of the substrate (the resistance of a material
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to changing temperature), the albedo of the surface (the fraction of incident solar energy in visual wavelengths that is reflected), and the surface slope (which determines the fraction of insolation received) [Putzig et al., 2005]. A higher thermal inertia material results in higher mean subsurface temperatures and hence the equilibrium condition for the ice table is met at a greater depth [Sizemore et al., 2009]. Significant localized subsurface temperature perturbations and ice table depth variations can be caused by changes in thermal inertia, for example, a decimeter scale rock can result in an ice free radius down to several decimetres depth [Sizemore and Mellon, 2006].
Figure 5-2: Depth of perennial ice table in the current martian climate, from models of mean surface temperature and thermal inertia, assuming zero elevation and water vapour density of 102 kg/m2 [Mellon et al., 2004b] (Fig. 3). The depth of the subsurface ice table on Mars under the present climate is located where the annual average partial pressure of water vapour in pore spaces is equal to the partial pressure of water vapour near the surface [Mellon and Jakosky, 1993]. At this depth, ice will be in diffusive equilibrium with the atmosphere with the net deposition of ice equal to the net sublimation on annual time scales. The diffusion of water vapour is dominated largely by the humidity gradient (which in turn is largely affected by temperature), with vapour diffusing from high to low partial pressure. The instability and depletion of mid-latitude shallow volatiles has likely been occurring since the last period of ~30˚ obliquity 10 Ma ago [Laskar,
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2004; Mellon and Jakosky, 1993], resulting in regions of desiccation. The depth of the current perennial H2O ice table from thermal models, indicating long-term ice stability within the regolith, is ~ 2 m at absolute latitudes of ~ 45˚ and decreasing with increasing latitude (Figure 5-2) [Mellon and Jakosky, 1993, 1995]. At low latitudes < 45˚, ground ice within at least 3 m depth is assumed to be unstable due to the warm soil temperatures frequently above 200 K within this layer. At the poles water ice can be stable on the surface due to the extremely low temperatures below the sublimation curve. Models of ice table depth are generally consistent with independent ice tables derived from subsurface hydrogen concentrations [Boynton et al., 2007; Mitrofanov et al., 2007] (5.2.3) and from groundtruthing. For example, at the Phoenix landing site in the northern arctic, ground ice was found at a mean depth of 4.6 cm under loose dry soil [Mellon et al., 2004b, 2008a], consistent with the equilibrium depth for vapour diffusion [Mellon et al., 2009a]. The ice was observed to accumulate and sublimate and re-accumulate within a day [Smith et al., 2009b], varying with the atmospheric Ppart-T. Ice that is not exposed to the atmosphere in this way may experience long term stability [Clifford and Parker, 2001]. This is possible when ice is isolated from the low atmospheric pressure by an impermeable strata such as regolith with a low porosity [Gilmore and Phillips, 2002] or ice-cemented pore space [Gaidos, 2001]. Sealing off the subsurface from contact with the atmosphere may also allow subsurface liquid to remain stable, if the ice table extends to the depth at which melting can occur.
5.1.2. Liquid The stability of liquid water has a similar dependence on Ppart. On Earth, the partial pressure of water vapour is much higher (point C, Figure 5-1) and only at high altitudes does it become low enough to restrict the presence of a liquid phase (2.4.1). On Mars however, the Ppart of water vapour in the atmosphere is always less than the triple point pressure (610 Pa) of pure water, and hence stability of liquid water on the surface is unlikely (red crosshatching, Figure 5-1). This does not mean that liquid cannot occur on the surface, but as it will never be stable its duration will be transient. Water stability is also determined by the total pressure Ptot as it affects the evaporation rate. Liquid water in contact with the martian atmosphere is unstable when both the total atmospheric pressure Ptot is < 610 Pa and Ppart < 610 Pa. However Ptot > 610 Pa occurs in some low altitude locations
105
such as the Hellas and Argyre basins [Carr, 2006] allowing liquid water to be metastable (red Mars circle, Figure 5-1). Metastability occurs when Ptot and T occur within the liquid regime, that is, the temperature is above the freezing point and Ptot > 610 Pa (for pure water) [Haberle et al., 2001; Richardson and Mischna, 2005]. The distinction between unstable and metastable is important, as it results in a different timescale for total loss of liquid water through evaporation or freezing. Any liquid water exposed to the martian atmosphere will be temporary, as the P-T environment restricts it to freezing and/or vapourising (depending on the location of Ppart, T within the phase diagram). In regions where liquid water is unstable it will vapourise rapidly [Carr, 2006]. However, the mean evaporation rate under metastable conditions, at T = 0 ˚C and Ptot = 700 Pa, is on the order of 0.7 mm/hr for liquid and 0.1-0.4 mm/hr for ice [Sears and Moore, 2005]. For example, if ~ 103 m3 of water carved a typical martian gully [Christensen, 2003] and pooled to a thickness of 1 m, it could have taken ~ 60 martian days to evaporate at 0 ˚C, 700 Pa. This is < 10 times faster than a typical annual average evaporation rate at sea level on Earth of ~0.1 mm/Earth hour5 [Penman, 1948]. Two factors serve to retard the evaporation and/or freezing of liquid water and lengthen its lifetime: water activity and atmospheric humidity. The water activity of an aqueous solution is the ratio of Ppart over the solution to Ppart over pure water at the same temperature [Blandamer et al., 2005] and is a measure of how freely the molecules can move and be available for reactions [Grant, 2004] (see 2.5). A solution with low water activity will have a lower P part over the solution at a given temperature, compared to a solution with higher water activity [Altheide et al., 2009]. This is due to the retardation of the energies of the water molecules at low activity [Koop, 2002], and hence a decrease in the fraction of water molecules with enough energy to enter the vapour phase.
The
evaporation rate is expressed in Equation 5-1, and the parameters are measured directly over the water surface (where the Ppart environment is largely determined by the solution). Although not explicit in Equation 5-1, evaporation will be higher in an unsaturated atmosphere, and low when the humidity is at 100 % and the 5
1 Earth hour ~ 1.005 Mars hours
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Literature Review & Background: Liquid Water on Mars
equilibrium Ppart occurs at its maximum value on the vapour curve, Psat (Figure 5-1) [Ingersoll, 1970]. In the atmosphere immediately over water on the martian surface the Ppart can be much larger than Psat for a short period of time, due to the energies of the surface water molecules and the lag in diffusion of water vapour [Chittenden et al., 2008]. Hence the relationship between Ppart and Ptot dominates evaporation and the rate increases with increasing Ppart and decreases with increased Ptot, as expressed in Equation 5-1.
Equation 5-1
ן ܧ
ೌೝ̴ೞೠ
̴ೌೞೝ
where E is the evaporation rate, and Ppart is measured over the solution at the ice/air or liquid/air interface [Chittenden et al., 2008; Kuznetz and Gan, 2002]. From Equation 5-1, the fastest evaporation would occur in an unsaturated atmosphere over a pure liquid water solution (high Ppart), at high altitude (low Ptot). The slowest evaporation would occur in a saturated atmosphere (Psat), over a low activity aqueous solution such as a brine (low Ppart) [Sears and Chittenden, 2005], at low altitude (high Ptot) [Altheide et al., 2009; Ingersoll, 1970]. The same is valid for the sublimation of water ice [Bryson et al., 2008]. If pressures are left unchanged, the evaporation rate is also retarded by lowering the temperature, as this retards the water activity (2.5.2) and lowers Psat. If liquid water is present on Mars it is likely to be saline, as a variety of salts have been detected in the martian regolith [Squyres et al., 2004a, 2004b] (5.3.3). Salts not only lower the freezing temperature of liquid water, but they also increase the lifetime of any liquid water that forms on the surface by lowering the humidity (RH) [Zorzano et al., 2009]. Modelling indicates that liquid water environments from melting of surface water frost could occur between 0˚-30˚ N in Amazonis Planitia, Arabia Terra, the Elysium region and in the southern hemisphere impact basins of Hellas and Argyre [Haberle et al., 2001; Richardson and Mischna, 2005] (low latitude frost discussed in 5.2.2). Fine grained materials assist in retarding the temperature and humidity variation in the subsurface, with the sublimation rate of ice under only 10 mm of fines half its surface value [Bryson et al., 2008]. In the aforementioned regions the conditions of warm surface temperatures, subsurface insulation by low thermal inertia materials, and high total pressure are
107
met. Liquid water may also form transiently on the surface from the melting of ice in high sun conditions at absolute latitudes ≥ 60˚ [Hecht, 2002] (5.3.4.a). Including salts expands the regions where liquid water can occur [Mohlmann, 2011], with modelling indicating that a concentration of 15-40 % salts (see 5.3.3) is sufficient to allow the melting of permafrost (or to maintain liquid water) in the top few metres of soil [Mellon and Phillips, 2001].
5.2.
Reservoirs of present day water ice
Locations of water ice are of astrobiological interest on Mars as they provide candidate environments for liquid water – either as thin films of liquid coating the surface of ice, or as meltwater generated under certain temperature and pressure conditions as discussed above. Given the transient nature of liquid water on the surface of Mars and hence the difficulty in observing it directly, the most effective strategy for searching for liquid is to identify where deposits of water ice intersect with P-T conditions that favour melting.
5.2.1. Polar caps Approximately 5-10 % of Mars’ water budget is thought to be stored as ice in the polar caps and circumpolar terrain. Water ice is present in the residual/perennial caps of both hemispheres [Langevin et al., 2007], exposed when springtime temperatures increase to the range of 150-205 K causing the overlying layer of seasonal CO2 ice to sublimate [Kieffer et al., 1976] while the underlying H2O ice remains stable [Jakosky and Farmer, 1982]. A small fraction of water frost is also exposed continuously at the edge of the northern seasonal cap [Kieffer and Titus, 2001; Wagstaff et al., 2008]. The northern residual cap is more extensive than the southern [Titus, 2005], with the southern water ice cap still retaining a thin covering several metres thick of CO2 ice even in the summer [Bibring et al., 2004; Montmessin et al., 2007; Titus et al., 2003]. This discrepancy between the two residual caps is caused by cooler regional weather patterns that move from the Hellas Basin to the polar region in the southern hemisphere [Colaprete et al., 2005; Giuranna et al., 2008].
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The orbital ground-penetrating radars- Mars Advanced Radar for Subsurface and Ionosphere Sounding (MARSIS) and Mars Shallow Radar Sounder (SHARAD)- have provided significant information on subsurface water ice on Mars [Mouginot et al., 2010; Phillips et al., 2011]. The difference between the dielectric constant of martian regolith and ice (~ 7.2 for basalt; 3.15 for solid H2O; 88 for liquid H2O [Beaty et al., 2001]) can be utilised by radar instruments to detect water, liquid and solid, as high dielectric losses result in a significant decrease in signal energy and hence a lower power received by the detector in orbit. A significant example of subsurface water ice detection is the polar layered deposits, which are composed of a mixture of dust, sand and ice with banded albedo [Carr, 2006]. The deposits, which are seen at both poles, are composed almost entirely of water ice identified through both hyperspectral observations [Bibring et al., 2005; Murchie et al., 2007] and a dielectric constant that is consistent with almost pure ice (from radar) [Picardi et al., 2005; Plaut et al., 2007]. These deposits, in addition to the residual water ice caps which they underlie, represent a significant volume of stored water as they extend for a thickness of ~ 3 km [Wieczorek, 2008]. The total volumes are dependent on estimates of the underlying topography of the cap [Zuber et al., 1998], but geometric interpolations provide an estimate of ~ 3×106 km3 for the combined total of water ice at both poles [Smith et al., 2001]. Potential subsurface melting of the polar water ice is discussed in section 5.3.5.
5.2.2. Surface frost The coverage of multi-spectral observations of Mars from the Compact Reconnaissance Imaging Spectrometer for Mars (CRISM) and the Visible and Infrared Mineralogical Mapping Spectrometer (OMEGA) now allows for highly localised and even seasonal deposits of surface water ice to be detected [Appéré et al., 2011; Brown et al., 2010].
Although small in volume and not spatially
extensive, these deposits still hold potential biological significance given that seasonal liquid water has been found to support psychrophiles in Earth’s permafrost and polar environments [Ponder et al., 2008; Price et al., 2002] (2.5). Seasonal surface water frost is common at latitudes above ~ 50˚ in both hemispheres (with some local variation [Mellon et al., 1997]). Equatorial water frost is more significant however, as the warmer average temperatures at low latitude allow for temperature conditions within shallow soil that are hospitable
109
for life and conducive to melting ice. The ground temperatures at the Opportunity landing site in late winter were > 190 K [Savijärvi and Kauhanen, 2008; Smith et al., 2006] and hence generally above the frost point. The warm surface temperatures and low daytime humidity at low latitudes negate widespread surface frost even in winter, due to rapid sublimation [Mellon et al., 2004b]. Particularly in equatorial regions, surface illumination conditions that provide local low temperature environments play a significant role in providing the thermodynamic conditions for temporary water ice [Nelli et al., 2009]. Slopes are an important factor, as they experience different conditions of solar insolation and temperature compared to a level surface [Mellon et al., 2000] and favour the cold-trapping of water vapour [Schorghofer and Edgett, 2006] (Figure 5-3).
Figure 5-3: A location in Terra Sirenum observed by CRISM, near 38.9 ˚S, 195.9 ˚E. CRISM bands cover 0.36-3.92 micrometres and show features as small as 18 m across. Blue indicates water ice which is deposited on the shadowed slopes. Image courtesy NASA/JPL-Caltech6. Recent spectral studies have identified that the effect of shadowing allows water frost to condense on some pole facing slopes throughout autumn to early spring at latitudes as low as 13 ˚S and 32 ˚N [Vincendon et al., 2010a], despite mean annual surface temperatures of > 200 K at these latitudes on flat surfaces [Mellon et al., 2004b]. The frost layers must be > 2 mm thick for detection, implying that they are
6
http://mars.jpl.nasa.gov/gallery/calibration/seasonalFrost.html
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Literature Review & Background: Liquid Water on Mars
accumulating over several days, with modelling indicating that deposits can grow to several centimetres thick. This detection of low latitude ice is consistent with an earlier publication of water frost on shadowed and pole-facing low latitude slopes throughout a similar range of seasons [Carrozzo et al., 2009], although the detection was complicated by the presence of ice clouds. Transient low latitude water ice environments may have astrobiological significance, as modelling suggests that unstable H2O ice can transition through the liquid phase prior to vapourisation [Kossacki and Markiewicz, 2004]. Depending on the volume and duration of seasonal liquid water derived from the melting of frost, it is plausible that it could percolate to hospitable temperature conditions within the shallow soil providing a potential seasonal hospitable environment for life. Furthermore, thin liquid films of mobile H2O molecules always occur in association with H2O ice [Haynes et al., 1992] (2.5.2) and would be expected to remain adsorbed to soil grains after the ice is lost [Mohlmann, 2004], although the ability of this form of liquid water to sustain active life is not clear [Jakosky et al., 2003].
5.2.3. Subsurface water ice The global occurrence of subsurface water-equivalent hydrogen on Mars was mapped by the neutron detector of the Gamma-Ray Spectrometer (GRS) suite and the Russian High-Energy Neutron Detector (HEND) onboard NASA’s Mars Odyssey. The orbital detection of gamma rays in combination with fast (> 500 kev), epithermal (0.4-500 kev) and thermal (< 0.4 ev) neutron detection allows the subsurface hydrogen (H) content within the top ~ 2 m of the martian subsurface (the penetration depth of cosmic rays) to be determined [Kuzmin et al., 2004]. Details on the instruments and data integration can be found in [Boynton et al., 1992, 2004] but the following provides a brief summary.
Cosmic rays
(predominantly protons and alpha particles) interact with the martian surface due to the thin atmosphere and the absence of a magnetic field and produce fast neutrons. These neutrons interact with other atomic nuclei by exciting them and causing them to decay to lower energies (thermal and epithermal) through the release of gamma rays.
Gamma rays in turn undergo scattering (with fast
neutrons) and capture (with epithermal neutrons) reactions, with the frequency of interactions being dependent on the elemental composition (hence neutron abundance) of the subsurface and the cosmic ray generated neutron flux. The
111
gamma rays received in orbit will not have the original unique emission energies produced by the de-excitation of each chemical element, as the gamma rays have lost energy through their neutron interactions. Therefore, the surface neutron flux must also be measured in order to deduce the elemental composition of the subsurface from the orbital gamma ray intensity. The neutron flux received in orbit is also strongly moderated by subsurface collisions. The most effective reduction of a neutron’s energy will occur in a collision with a H nucleus due to its comparable mass (1 proton).
Hence a
concentration of H atoms in the subsurface will reduce the flux of energetic neutrons (fast and epithermal) and increase the flux of thermal neutrons escaping the surface. Neutrons at thermal energies can also be captured by some nuclei however. Hence, knowledge of the capturing and scattering cross sections of various chemical elements is required to interpret the flux of the three energy bands of neutrons received at the detector. Ultimately, monte-carlo simulations are utilised to calculate the expected neutron and gamma ray flux for various models of regolith composition. Although most sensitive to hydrogen content, maps of other elemental abundances can also be produced [Boynton et al., 2007]. The general trends observed for subsurface H are: (i) decrease in high energy epithermal and fast neutrons with increasing H; (ii) increase in thermal neutrons for increasing H up to 10 wt%; (iii) decrease in thermal neutrons for increasing H above 10 % and increasing burial depth of H layer. The estimated accuracy of derived hydrogen concentrations through this method is ±2 % [Feldman et al., 2011].
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Figure 5-4: Water-equivalent hydrogen concentrations within the top 2 m depth of the martian subsurface from orbital neutron flux [Maurice et al., 2011] (Fig. 32). Pixel size is ~ 600×600 km.
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Hydrogen concentrations detected by the method above are shown in Figure 5-4. At high latitudes the hydrogen has been interpreted as H2O ice [Feldman et al., 2002; Mitrofanov et al., 2002]. The GRS/HEND results indicate that the martian permafrost has an increasing water ice content polewards of absolute latitudes ≥ 50˚ [Mitrofanov et al., 2007] with the concentration varying between 13 wt% (~ 50˚ absolute latitude) to ice-rich with 80 wt% (~ 90˚ absolute latitude). This is generally consistent with the surface spectral observations of 3 µm H2O absorption at 5-15 wt% polewards of 45˚ [Milliken et al., 2007]. The water-rich substrate is buried beneath a desiccated layer of soil that is ~ 10’s of centimetres thick at ~ 50˚ to only millimetres thick or absent at the poles [Feldman et al., 2002; Prettyman, 2004]. The global hydrogen abundance is moderately correlated with chlorine (partial correlation coefficient between 0.3 [Karunatillake et al., 2006] and 0.2 [Keller et al., 2006]). The hydrogen distribution is still poorly explained however, as only 20% of its variance is accounted for by Cl and it shows no strong trends with other elemental abundances, mineralogy, or physical parameters such as thermal inertia, rock abundance and albedo [Karunatillake et al., 2006]. This can be in part explained by the significantly different spatial resolution between the hydrogen map (~ 600×600 km pixels) and the thermal inertia and albedo maps (~3×3 km pixels) [Boynton et al., 2002b; Maurice et al., 2011]. The link between H and Cl suggest both elements moved together [Boynton et al., 2007; Keller et al., 2006], most likely through a weathering related process [Hahn et al., 2007]. The increasing gradient in H concentration from the mid-latitudes towards the poles clearly has a correlation with decreasing insolation and annual average surface temperatures, and is consistent with models of the depth to perennially stable water ice (5.1.1). In the lower latitudes ≤ 50˚, high mean surface temperatures result in shallow ice being unstable. The hydrogen content at these latitudes was originally attributed to OH within hydrated minerals (the least volatile form of H2O) [Feldman, 2004; Mellon et al., 2004b]. The H at these latitudes is significantly lower, varying predominantly between 1-10 wt%. The upper fraction occurs in several enriched regions in Arabia Terra and Medusae Fossae [Mitrofanov et al., 2003]. The most depleted regions with ~ 0.24 wt% [Boynton et al., 2002b] occur in Solis Planum and western Argyre [Litvak et al., 2006]. These low latitude concentrations are
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generally consistent with the spectral observations of 3 µm H2O absorption at < 5 wt% equatorwards of 45˚ [Milliken et al., 2007]. The low hydration of the soil can be interpreted as either H2O or OH that is chemically (within a hydrated mineral matrix) or physically (on the surface of grains, as in thin films) bound [Mitrofanov et al., 2007]. As only 0.1 wt% of water is expected to coat shallow subsurface grains from vapour condensation [Mohlmann, 2002, 2005], the low latitude H content is more likely bound water in the form of hydrated minerals [Mitrofanov et al., 2002]. Even under the low partial pressure of water vapour on present day Mars, there are mineral formations, such as clays and sulphates, that are able to capture atmospheric water within their matrix and release it on diurnal time scales, providing a source of water to the subsurface [Bish et al., 2003; Fialips et al., 2005; Janchen et al., 2006]. OH within minerals (such as hematite) is unlikely to be unavailable however [Yen et al., 1999], as significant heating is required to liberate water molecules from many mineral matrices [Basilevsky et al., 2003]. Consistent with this interpretation of the low hydrogen concentrations, spectral observations of OH/H2O in the infrared suggest that the low latitude H may be partially due to hydrated sulphates, phyllosilicates and iron-oxides within martian dust [Pommerol et al., 2011]. Recent evaluation of GRS results using sophisticated models of subsurface layering suggest that some of the enriched low latitude hydrogen may be attributed to buried water ice in regions such as Arcadia Planitia and Promethei Terra [Feldman et al., 2011]. A further alternate explanation for the < 10 wt% concentrations of subsurface H is transient, unstable ground ice [Jakosky et al., 2005]. This latter hypothesis would be testable by comparing the deposition of water surface frost in an enriched low latitude region such as Arabia Terra, with an area of H depletion.
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Figure 5-5: Lobate debris apron in Niger Vallis, an ancient channel feeding into Hellas Planitia at 35.6 ˚S, 268 ˚W. Feature is likely an ancient glacier covered by rocky regolith. HiRISE image courtesy NASA/JPL/University of Arizona7, image width ~ 280 m. A number of other methods have been employed to investigate subsurface water ice and their results should be briefly mentioned. As discussed in 5.2.1, radar has been utilised previously to detect the subsurface extent of the martian polar caps and polar layered terrain. More recently, it has been employed to detect metres thick accumulations of water ice [Holt et al., 2008; Parsons et al., 2011] associated with the glacial morphologies termed lobate debris aprons [Squyres, 1978] (Figure 5-5). These features occur regionally, with a substantial concentration observed east of the Hellas Basin, and may reflect relatively young (< 500 Ma) ice remnants from higher obliquity that provide an important tracer of martian climate change [Fastook et al., 2011; Morgan et al., 2011]. With the increase in sophistication and accuracy of general climate or circulation models for Mars [Forget et al., 1999; Hollingsworth et al., 2008], discrepancies between modelled conditions and actual 7
http://hirise.lpl.arizona.edu/PSP_002123_1440
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observations can be used to infer conditions in the subsurface. For example, some authors postulate that the lack of observable carbon dioxide frost on slopes where the atmospheric conditions indicate it is stable is evidence for an increased subsurface thermal inertia in those regions [Vincendon et al., 2010b] (enabling higher subsurface temperatures), attributed to the presence of a layer of water ice. This avenue of logic is however predicated on the completeness of the climatic model in explaining the surface, atmosphere, and their interactions, which is unlikely given the complexity of the martian systems being modelled. Finally, the presence of deep subsurface ice has been uncovered through excavation by a number of new mid-latitude impact craters [Byrne et al., 2009], the majority of which occurred in the enriched hydrogen region of Arcadia Planitia discussed above. The conditions under which this ice sublimated indicate that it is nearly pure water ice.
5.2.4. Surface morphologies indicating subsurface ice A great deal has been learnt about Mars through drawing analogies to terrain that is observed and understood on Earth.
Martian polar processes such as the
formation of permafrost, ice lakes [Warner et al., 2010] and ice-rich morphologies such as glaciers [Soare et al., 2012], pingoes and polygonal terrain are all present on the Earth in regions such as the Antarctic Dry Valleys, Iceland and the Arctic [Farr, 2004]. Several of the most compelling indicators of martian surface features that may have formed through the involvement of significant concentrations of ice will be summarised here. The current methods of direct detection of water ice on Mars are limited by low spatial resolution (such as the water-equivalent hydrogen map) or low sampling (such as spectral signatures of water and radar observations), and hence the surveying of morphologies indicative of subsurface ice can greatly assist in the search for water environments.
5.2.4.a.
Rampart craters
Rampart craters have a distinct morphology not seen on the Moon or Mercury, with a lobate rather than circular debris apron featuring an elevated edge and often fluvial morphologies throughout the ejecta blanket [Mouginis-Mark, 1979] (Figure 5-6). The dominant theory for rampart crater formation is excavation of
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the impactor into material with a high volatiles content, which when melted by the impact heating causes the ejecta to flow across the surface [Baloga et al., 2005; Costard, 1989; Mouginis-Mark and Baloga, 2006; Mouginis-Mark, 1987; Mutch and Woronow, 1980]. This interpretation is based on laboratory and model data (such as viscosity and surface flow calculations) which indicate that fluids were present in the ejecta of rampart craters and hence ice was present in the subsurface.
Figure 5-6: Comparison between martian rampart craters Yuty and Wohoo displaying fluidized ejecta (left; 22.2 ˚N, 325.9 ˚E), and lunar crater Timocharis with dry rayed ejecta (right; 26.7 ˚N, 13.7 ˚W). On the left, the composite of Thermal Emission Imaging System (THEMIS) images was constructed using Java Mission Planning and Analysis for Remote Sensing (JMARS)8. Right image taken by Apollo 15, image courtesy NASA/JPL-Caltech9. For scale, diameters of Yuty and Timocharis craters are ~ 18 km and ~ 34 km respectively. The best candidate ejecta fluid on Mars is liquid water from the impact melting of water ice [Carr et al., 1977; Gault and Greeley, 1978]. Currently, there is no spectral evidence of H2O ice, CO2 ice, or minerals that would be expected from water-rock interaction within rampart crater ejecta, however there has not been a THEMIS images: I11073008, I08215018, I03359002, I02660002, I01886009, I01524004. I00850009, V10449014, V14742009; metadata in Appendix A, Table A2; processing details in Appendix B. 9 http://photojournal.jpl.nasa.gov/ 8
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significant survey to look for such.
Rampart craters are strongly spatially
correlated with past water on Mars – being frequently associated with other morphological indicators of water (such as giant polygons [Martinez-Alonso et al., 2011] discussed in 5.2.4.b, ancient spring structures [Allen and Oehler, 2008] and outflow channels [Jones et al., 2011]). Rampart craters are also correlated with moderate to high present day subsurface hydrogen concentrations [Dohm et al., 2007], indicating that water ice may still be present within or beneath their ejecta (depending on whether ejecta thickness exceeds the 2 m hydrogen sensing depth). For example, potential water ice has been recently observed in an ejecta blanket through shallow radar [Nunes et al., 2011]. Furthermore, ice has been observed infilling the floor of some rampart craters [Hsu et al., 2009]. Hence some rampart craters may be locations of present day shallow water ice. Based on the assumptions that: (1) the preferential occurrence of rampart craters in some regions indicates that abundant volatiles were present in the subsurface in the past; and (2) that the excavation depth of a rampart crater reflects the depth of the ice-rich zone at the time of impact [Kuzmin, 1980]; then rampart craters can be used to identify regions of shallow ice abundance. When selected for terrain age, rampart craters can be utilised to trace the depth of water ice concentrations through martian history [Barlow and Bradley, 1990; Barlow, 2001; Barlow et al., 2001; Reiss et al., 2005a, 2006a]. Studies of regional trends in rampart crater onset diameter (the smallest diameter, and hence shallowest crater that is identified as a rampart) have located several equatorial regions that may have been ice rich in the relatively recent martian past. For example, within the Amazonian aged (3.0Gapresent) Chryse Planitia rampart craters have an average diameter of ~ 4 km with onset diameters indicating shallow ice from the past 100 Ma within the top ~ 60 m at the time of impact (Figure 5-6) [Costard, 1989; Demura and Kurita, 1998; Reiss et al., 2005b]. These geologically young craters, and any fresh rampart craters that are observed to form, can relate important information on the depth and concentration of the ice table under current climatic conditions [Osinski et al., 2011]. Rampart craters are particuarly useful as they can provide information on subsurface ice beneath the maximum depth sensitivity of the GRS (~2 m), SHARAD (~500 m) and MARSIS (~4 km) instruments [Fois et al., 2007]. For example, the rampart crater Yuty (Figure 5-6) has a depth of ~ 750 m (from MOLA datum).
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More generally, using the depth-versus-diameter relationship for ramparts craters [Reiss et al., 2006b], crater with diameters approximately 3.8 km and greater may have excavated to depths exceeding
500 m.
Furthermore they are strong
astrobiological targets [Beaty et al., 2006] as impacts into ice-cemented regolith may have triggered subsurface melting, fluid flow within rock fractures, and temporary hydrothermal environments.
5.2.4.b.
Polygons, pingoes
Polygonal fractures in the martian terrain occur widely at high latitudes above 55˚ in both hemispheres. Morphologically similar features occur on Earth, termed icewedge polygons, and form from repeated fracturing of permafrost by thermal contraction [Burr et al., 2009]. Cracks formed fill with meltwater in the spring and summer which then freezes within the cracks as temperatures cool, widening them [Black, 1976]. This mechanism is a poor fit to some martian examples however, as the larger scale of martian polygons (up to 20 km across) makes this mechanism unfeasible [McGill, 1986]. Although thermal contraction may be responsible in forming polygons 10-300 m across [Mangold, 2005], larger polygons may be formed from alternate processes involving subsurface ice, such as slumping or warping of the terrain after ice is lost through sublimation from the pore space, or from tectonic uplift [Hiesinger and Head, 2000]. The most important distinction between the two aforementioned mechanisms is the action of liquid water as the formation of small polygons generally requires some presence of episodic liquid water at the surface [Seibert and Kargel, 2001]. Hence, young small polygons would be significant, as they could provide clues to geologically recent liquid water. Such small polygons are difficult to date however, as they occur on terrain from each geologic epoch but have too few craters superimposed to enable absolute ages to be derived [Levy et al., 2009]. Polygon ages fall within the broad range of 0.1-100 Ma [Levy et al., 2009; Mangold, 2005]. The link between polygons and water ice is strengthened by their association with other putative ice morphologies (such as pingoes) and with the water-equivalent hydrogen map (5.2.3). Pingoes are circular elongated mounds that form on Earth from the freezing of subsurface water that forces the ground upwards through hydrostatic pressure [Burr et al., 2009; Gurney, 1998]. One extensive location of
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both polygon and pingoe morphologies occurs in Utopia Planitia [Depablo and Komatsu, 2009; Soare et al., 2005]. This region is believed to have been covered with water during the Hesperian (surfaces aged 3.7-3.0 Ga) and displays many fluvial morphologies [Fairen and Dohm, 2004; Tanaka et al., 2005]. Furthermore, due to its proximity to the Elysium volcanic region it likely experienced higher heat flow [McGovern, 2002] and possible hydrothermal activity.
Hence the past
presence of ponded water within the region, which would have frozen as the climate shifted to lower obliquity and the high latitude surface temperatures decreased ~ 3 Ga ago [Warner et al., 2010], supports the interpretation of polygons and pingoes as formed by water ice driven processes.
Globally, these
morphologies are correlated with regions of enriched hydrogen in the top ~2 m of soil above 55˚ [Mangold, 2004].
This depth is consistent with the depth
propagation of seasonal temperature variations required to initiate thermal contraction of ice (discussed above). Identification of geologically recent volatiles, such as in Athabasca Valles and Marte Vallis [Balme and Gallagher, 2009; Page, 2007], suggests the possibility of a subsurface reservoir of ice potentially extending to the depth of the melting isotherm and generating liquid water beneath [Clifford, 1993]. In general however, these features have more potential to inform on the distribution of past surface volatiles and climatic variation [Levy et al., 2010; Soare et al., 2011] than they do to identify regions of present-day shallow ice and possible liquid water.
5.3.
Present day liquid water
Despite being small in volume and transient in nature, liquid water may be a frequent constituent in martian surface environments. Below, the signatures of present-day liquid water at or near the surface of Mars will be discussed and correlated with locations of water ice.
5.3.1. Deliquescence & water adsorption Deliquescence is the phenomenon responsible for one of the most compelling potential observations of liquid water to date on Mars. Deliquescence is a phase transition from a dry solid to a droplet of saturated solution, which occurs when a particle absorbs water from the atmosphere onto its crystalline surface until it is
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liquefied and dissolved [Tung, 2009].
The process is driven by changes in
humidity, with the size of the water droplet increasing with increasing water vapour pressure [Lewis and Schwartz, 2004]. There is a delicate balance, however, between the volume of water and the concentration of solute required to decrease the freezing point sufficiently to allow it to remain a liquid. If the liquid volume is too large and the concentration of the solute becomes too low, freezing and sublimation will occur. Hence there is a maximum size to which water droplets can grow on the martian surface, dependent on the availability of the solute and the atmospheric Ppart and T. The mechanism of deliquescence has been proposed as a plausible explanation for the spherical droplets observed on the struts of the Phoenix Lander (Figure 5-7) in the northern martian arctic (68 ˚N), with perchlorate salts as a candidate solute [Rennó et al., 2009]. The perchlorate anion, ClO4-, typically bonds with a cation such as Mg+ or Na+ to form a salt, and was detected by Phoenix in concentrations of 0.4-0.6 wt% in the soil [Hecht et al., 2009]. Optimally a concentrated solution of > 50 % perchlorate salts [Smith et al., 2009b] could depress the freezing point of water by as much as 75 K (to 198 K) [Besley and Bottomley, 1969; Marion et al., 2009a], however under martian atmospheric conditions the salts form an aqueous solution down to a temperature of 205 K [Zorzano et al., 2009]. The observations at the Phoenix landing site provided valuable information on the likely state of surface or near-surface liquid water on Mars. Small volumes of brine coating soil and ice grains are most likely the dominant form of shallow liquid water on Mars [Mohlmann, 2008]. Deliquescence may occur anywhere that salts with a low eutectic temperature (the maximum value of freezing point depression) are present and the relative humidity is above a certain threshold value [Rennó et al., 2009]. The ability of salts on Mars to temporarily liquefy by the absorption of atmospheric liquid water provides an important source of water for potential biology within the subsurface diffusion depth of water vapour. Furthermore, the significant freezing point depression caused by perchlorate salts could be responsible for the generation of deep subsurface brine, as assuming a 1% perchlorate salt component in other martian soils could induce basal melting of the northern polar cap [Fisher et al., 2009, 2010].
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Figure 5-7: Spheroids on the leg of the Phoenix Lander, photographed by the robotic arm camera [Rennó et al., 2009] (Fig. 3). Spheroids have been interpreted as liquid perchlorate brines due to their liquid-like behaviour of spheroid growth with the peak spheroid size observed during high humidity in the afternoon. Large spheroids were observed to detach from the lander leg.
Furthermore, the presence of
perchlorates in the martian soil, which have a eutectic temperature consistent with remaining a liquid throughout the temperature range of the lander legs during the lifetime of the droplets, provides additional evidence for the droplets being a liquid brine [Hecht et al., 2009; Rennó et al., 2009].
5.3.2. Thin films More generally, low volumes of surface liquid water can occur without the presence of salts. Instead, liquid can become adsorbed by a surface (adhesively bonded) as a thin film (described in 2.5.2). On Mars, these films would be sourced from the condensation of water vapour below 200 K [Mellon and Phillips, 2001] (the frost point, section 5.1) to coat soil grains, or would occur as a coating on water ice or frost in shallow soil [Anderson, 1967] (analogous to unfrozen water in Earth permafrost [Anderson and Tice, 1972]). Thin films of pure or briny liquid water less than 1-2 water molecules thick (< 5 Å) at temperatures below -20 ˚C (153 K) [Mohlmann, 2005], can persist to temperatures of -100 ˚C (173 K) [Pearson and Derbyshire, 1974] (or potentially as low as -130 ˚C [Mazur, 1980]). They are likely to be found through the soil wherever it is in contact with a source of water. As with all liquid water exposed to the atmosphere, high humidity serves to
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stabilise the liquid by retarding its rate of evaporation [Altheide et al., 2009], thus lengthening its lifetime in the martian environment (5.1.2). Table 5-1: Eutectics of saturated aqueous solutions of Mars candidate salts. Salt*
Eutectic ˚C
Reference: freezing point depression
Na2SO4
-5
CaSO4
-0.06**
[Post, 1977; Williams and Robinson, 1981] [Marion and Farren, 1999]
MgSO4
-5
FeSO4
-68
[Pillay et al., 2005; Zhang and Chan, 2000] [Chevrier et al., 2009]
K2SO4
-1.6
[Marion and Farren, 1999]
NaCl
-21
[Grant, 2004; Greenspan, 1977]
CaCl2
-54
MgCl2
-33
[Bui et al., 2003; Morillon and Debeaufort, 1999; Siegel et al., 1979; Siegel and Roberts, 1966] [Greenspan, 1977; Morillon and Debeaufort, 1999; Siegel and Roberts, 1966]
CaCO3 MgCO3
-1.9 -0.3
[Marion, 2001] [Marion, 2001]
FeCO3 NaNO3
-18
[Greenspan, 1977; van der Ham et al., 1998]
CaNO3 Mg(NO3)2
-32
KNO2
-40.2
NaBr
-29
MgBr2
-42.7
[Greenspan, 1977; Morillon and Debeaufort, 1999] [Morillon and Debeaufort, 1999] [Greenspan, 1977; Morillon and Debeaufort, 1999; Negi and Anand, 2004; Tang et al., 2003] [Morillon and Debeaufort, 1999] [Besley and Bottomley, 1969; Zorzano et al., 2009]
(2Mg,Na)ClO4 -69 perchlorate *Grey **Has
shading indicates values not found in the literature. not been tested experimentally.
Reference: observed/hypothesised presence on Mars [Altheide et al., 2009; Arvidson et al., 2010] [Altheide et al., 2009; Bishop et al., 2005] [Altheide et al., 2009; Bishop et al., 2005] [Bishop et al., 2005; Chevrier et al., 2009] [Clark et al., 2005; Marion et al., 2009b] [Altheide et al., 2009; Murchie et al., 2009; Osterloo et al., 2008; Wray et al., 2009] [Altheide et al., 2009; Murchie et al., 2009; Osterloo et al., 2008; Wray et al., 2009] [Altheide et al., 2009; Murchie et al., 2009; Osterloo et al., 2008; Wray et al., 2009] [Boynton et al., 2009] [Ehlmann et al., 2008; Palomba et al., 2009] [Morris et al., 2010] [Manning et al., 2008] [Manning et al., 2008] [Manning et al., 2008] [Manning et al., 2008] [Marion et al., 2009b]
[Marion et al., 2009b] [Hecht et al., 2009]
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5.3.3. Brines In the subsurface, the presence and salinity of liquid water can only be speculated upon, but at least temporary liquid water can be expected wherever there is a source of water ice and temperatures above the eutectic. A number of salts that depress the freezing point of water are present on the martian surface, and are expected throughout the crust (Table 5-1). The presence of soluble compounds in the martian soil suggests that saline solutions may exist in the martian subsurface [Clark and van Hart, 1981]. The current knowledge of the salts present on Mars allows an informed estimate to be made of the likely salinity and therefore the range of possible temperatures of any liquid water in the martian soil. Sampling by the surface missions, as well as remote sensing by OMEGA and CRISM, has revealed concentrations of salts in martian terrains. The observed salts are Mg, Ca, K and Na sulfates [Gendrin et al., 2005; McLennan et al., 2005; Squyres et al., 2004a], Ca, Fe and Mg carbonates [Boynton et al., 2009; Ehlmann et al., 2008], and chlorides (most likely Na, Mg and/or Ca [Clark et al., 2005; McLennan et al., 2005]). These and other potential martian salts are summarised in Table 5-1.
Salts
promote the occurrence and duration of surface brines through deliquescence, or subsurface brines if liquid water is available. Sulphates have been sampled at the Pathfinder, Viking and Opportunity landing sites at concentrations between 6-40 wt% in the soil [Foley, 2003; Moore, 2004], and remote sensing spectroscopy indicates that in some locations the fraction of magnesium sulphate may reach 30 wt% [Vaniman et al., 2004]. Carbonates have been detected from orbital spectra and sampled directly, and appear to be globally widespread with concentrations varying from: 2-5 % to 15 wt% from remote sensing [Palomba et al., 2009]; and 334 % from surface sampling by Spirit and Phoenix [Boynton et al., 2009; Morris et al., 2010]. Chlorine in some form is globally distributed on the martian surface at ~ 0.2-0.8 wt% [Boynton et al., 2007] has been sampled in the soil at < 1 wt% [Bruckner et al., 2011; Rieder et al., 2004]. The cation in chloride salts is difficult to identify, as chlorides lack diagnostic features in the visible and near infrared [Murchie et al., 2009], however the dominant cations are expected to be Na, Mg and Ca [Kuzmin and Zabalueva, 2000; Marion et al., 2009b]. Nitrates have not been detected on Mars [Jakosky, 2007], despite the ability of CRISM to detect them [Murchie et al., 2007]. This lack of detection will become increasingly significant as
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the coverage of CRISM on the surface increases, however given that < 1 % of the martian surface will be observed by CRISM at high-resolution [Murchie et al., 2006], in-situ detection remains the best strategy.
The composition of a
subsurface martian brine would vary locally but, given the observed abundances of salts and their respective solubilities [King et al., 2004], the dominant constituents in an aqueous solution can be estimated. In synthesis, subsurface brines on Mars might be composed primarily of magnesium sulfate, calcium sulfate (total sulphates < 8 % [Marion et al., 2010]), sodium chloride, magnesium carbonate and calcium carbonate [Boynton et al., 2009; Jagoutz, 2006; Marion et al., 2009b; Tosca et al., 2008]. Measurements of martian soil indicate that the total salt varies typically between 8-30 % by mass (Viking [Clark and van Hart, 1981]; Pathfinder [Wänke et al., 2001]; Spirit and Opportunity [Vaniman et al., 2004]; Phoenix [Hecht et al., 2009]).
Given the
information in Table 5-1, most candidate salts on Mars have eutectic freezing temperatures in the range of -10 ˚C to -20 ˚C. The few exceptions to this are salts which have not been detected on Mars – such as FeSO4, CaCl2, and bromides- and perchlorates which are unlikely to occur in substantial concentrations within the soil or a putative brine. Hence it is conservatively estimated that an average brine on Mars would have a weakly depressed freezing point around 263 K from sulphates and chlorides. The freezing point will depend on the number of ions in the system and the complex equilibrium relationship between multiple phases of fluid–solid–vapour [Jagoutz, 2006]. Subsurface liquid water on Mars that is not in contact with the atmosphere could potentially be highly saline as it may not be recharged by fresh liquid water or water vapour. The phase diagram of this putative martian brine will be modelled in Chapter 6. Although there is currently no direct detection of a martian brine, the suggestive evidence of their existence will be discussed below.
5.3.4. Surface fluvial morphologies Water in the liquid phase has not been definitively detected on Mars (although controversial spectral evidence of brine was recently announced [Rennó and Mehta, 2011]), however an increasing number of fluvial morphologies are being
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documented and monitored for episodic activity and spectral signatures of water. The most compelling examples of these surface features will be summarised here.
5.3.4.a.
Gullies
Martian gullies present the best suggestive evidence for recent and potentially present day [Malin et al., 2006] liquid water on Mars. Gullies on Mars are so named as they are morphologically similar to Earth gullies, however their large scale and the method of their formation (rainfall does not occur on Mars) makes them quite distinct from the majority of the Earth counterparts. Martian gullies are generally recognised by the following features: an alcove that forms on a slope representing where material has been eroded away and can measure several hundred metres across [Christensen, 2003]; a channel emanating from the base of the alcove and ending downslope that can be several kilometres long [Heldmann and Mellon, 2004]; and a debris apron at the termination of the channel interpreted as the site of deposition of the material transported by the flow [Malin et al., 2006] (Figure 5-8). Cold climate gullies on Earth display the same structure and are formed through the action of liquid water [Coleman et al., 2009; Dickson and Head, 2011], and hence they provide qualitative evidence for the involvement of a fluid in the formation of the martian analogue. The visual similarities between the features on Earth and Mars however do not exclude the possibility that martian gullies could be formed by dry gravity driven debris flows down slope, lubricated by at most a small percentage of soil moisture [Bart, 2007; Treiman, 2003], or by liquid carbon dioxide [Hoffman, 2002].
This review will continue under the
assumption that they form from liquid water as this is the most likely hypothesis for the majority of gullies, and examine what information they can potentially provide on recent liquid water on Mars.
It is worth noting that numerical
modelling (discussed below), orientation preference for pole-facing slopes [Dickson et al., 2007], channel behaviour (such as redirection around obstacles and influence of surface relief [Conway et al., 2011a]) and a strong correlation of gully occurrence with high solar insolation [Morgan et al., 2010] and high thermal conductivity regolith [Heldmann et al., 2005], indicate that many of the gully systems on Mars could be fluvial in origin [Balme et al., 2006; Kneissl et al., 2010].
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Figure 5-8: Gullies in Corozal crater in Terra Cirrenum at 159.4 ˚E, 38.7 ˚S (HiRISE image10). There is great variety in channel thickness and degradation of the gully systems. Some gullies appear fresh and have bright deposits. Boulders can be seen in and around the gully alcoves. Gullies predominantly form on the pole-facing slopes. Red square gives the extent of the top image on the right, green square of the bottom image. As martian gullies are strongly suggestive of fluid being present and at least temporarily stable on the surface, it is extremely important to constrain the timeframe and hence the orbital and climatic conditions under which they formed. The upper limit of gully ages is 1.25 Ma [Reiss et al., 2004; Schon et al., 2009] which 10
HiRISE image ESP_014093_1410; metadata in Appendix A, Table A1; processing details in Appendix B.
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places their formation as subsequent to the last period of obliquity higher than 30˚ [Laskar, 2004], during which time extensive ground ice metres thick was deposited in the mid-latitudes where most gullies occur [Costard et al., 2002; Head et al., 2003]. It is unlikely that significant amounts of this ice pack remain today [Farrell et al., 2009; Nunes et al., 2010; Williams et al., 2008]. The re-activation of two old gullies within a ten year period [Malin et al., 2006] and the seasonal activity of a number of gully systems (discussed below) indicates however that the processes that form them and drive fluid to the surface are currently active, although localised and highly sporadic. This makes them strongly relevant to the search for current liquid water environments and potential microbial habitats on Mars. One theory for gully formation is by liquid water from a subsurface aquifer. In this model, the gully fluid exits explosively from beneath the surface (due to the pressure gradient between the aquifer and the atmosphere) and the depth of the gully alcove beneath the local surface is indicative of the depth of the source fluid. This depth varies between 0- 2 km beneath the local surface with a global average of ~ 200 m [Heldmann and Mellon, 2004; Heldmann et al., 2007].
Thermal
modelling indicates that pure liquid water is possible at the majority of studied gully depths as the modelled temperatures are above melting if a layer of dry soil at the surface insulates icy soil beneath [Heldmann and Mellon, 2004; Heldmann et al., 2005]. This model of icy soil beneath ~ 10’s of centimetres of dry soil is consistent with the hydrogen results for mid-latitudes [Mitrofanov et al., 2004]. Varying thermal conductivities and invoking salts to lower the freezing point of water [Knauth and Burt, 2002] allows water to be liquid at the remaining studied gully depths [Chevrier et al., 2009]. The association of gully systems with a distinct rock layer on a crater wall [Mellon and Phillips, 2001], indicative of impermeable strata [Gilmore and Phillips, 2002], is consistent with this theory as such a layer is required to confine liquid water and stabilise it under current climatic conditions. Furthermore, modelling indicates that multiple aquifer discharge events over periods of months are consistent with the gully features [Goldspiel and Squyres, 2011; Marquez et al., 2005; Parsons and Nimmo, 2010], with the aquifer plausibly fed by the melting of remnant subsurface snowpack from higher obliquity [Head et al., 2008]. This is supporting evidence for subsurface liquid in the recent martian past, which may still persist in association with some gully systems. It is worth
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noting, however, that there is no current evidence for deep aquifers as the deepest water ice detected at 4 km depth (5.2.3) is much shallower than the likely depth of the melting isotherm [Clifford, 1993]. An alternate theory is the formation of gullies through meltwater produced seasonally at much shallower depths than the average ~ 200 m alcove depth. Potential sources for this water are water rich snowpack deposited at midlatitudes during former higher obliquity [Christensen, 2003; Schon and Head, 2011a], surface frost cold-trapped into the gully alcove and producing thin flows of liquid as it sublimates [Dickson et al., 2007] or seasonal melting of permafrost [Conway et al., 2011a; Lanza et al., 2010; Levy et al., 2011]. In these scenarios, liquid water was generated transiently at the surface and down to depths of ~ 20 cm within a snow or ice layer [Bridges and Lackner, 2006; Head et al., 2008; Mellon and Phillips, 2001]. Gullies occur at mid- to high latitudes of both hemispheres poleward of 30˚ and are correlated with both the locations of snowpack deposition during higher obliquity and modelled conditions of melting under past climate conditions [Williams et al., 2008, 2009]. Although snowmelt is unlikely to feed most gullies under modern martian obliquities [Kossacki and Markiewicz, 2004; Williams et al., 2009], it is worth noting that gullies are weakly correlated with current subsurface hydrogen content > 10 %, which is interpreted as water ice filling pore space [Boynton et al., 2002b], and it is plausible that small amounts of shallow liquid are still generated through solar heating at some gully sites [Marquez et al., 2005].
Modelling indicates that favourable conditions for
meltwater formation from atmospherically deposited surface frost can still occur in some mid-latitude gullies [Heldmann et al., 2005; Kossacki and Markiewicz, 2004]. As gully alcoves are sheltered, often in shadow due to their slope and orientation, they are ideal for cold-trapping water vapour and may frequently support transient liquid water environments (5.1.2).
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Figure 5-9: North polar gullies at 68.5 ˚S, 1.3˚ E (HiRISE image11). Ice is coating the gully alcoves, channels and aprons. Recent flow was observed through some gullies in this system [Diniega et al., 2010]. Numerous recent gully activity has been observed however the observations provide limited clarity to the questions of how gullies formed and where the source fluid was located. Although studies of new deposits (Figure 5-9 and Figure 5-10) found a seasonal preference there are discrepancies between them (Table 5-2). Some new flows occurred during winter and autumn, suggesting that temperature may be too low for liquid water [Dundas et al., 2010]. Others
11
HiRISE image ESP_011396_1115; metadata in Appendix A, Table A1; processing details in Appendix B.
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however occurred during autumn and spring and are potentially consistent with liquid water [Harrison et al., 2009]. In general deposits were originally bright [Malin et al., 2006; McEwen et al., 2007a] and darkened over time, but no evidence of salts or water was identified [Barnouin-Jha, O. S. McGovern et al., 2008]. None of the observations found frost in association with the gullies, nor provided temperatures of the surface materials at the time of the flows to determine whether they are above the frost point. It is therefore not clear whether new flows are associated with volatiles (H2O or CO2) or whether they are avalanches of dry material [Kolb et al., 2010]. If liquid water was involved it was low volume, and hence it is likely that the new flows were fine-grained debris lubricated by thin films of liquid water sourced from frost or atmospheric water vapour (such as Atacama gullies [Heldmann et al., 2008]). Methods such as comparing the slope of the crater wall to the length of the gully channel may be able to distinguish between some scenarious of ‘dry’ and ‘wet’ formation [Capitan et al., 2011]. Furthermore, it is also possible to envision an evolution of the gully formation mechanism through time [Capitan et al., 2011; Schon and Head, 2011b], with ‘wet’ formation being dominant when mid-latitude ice was being emplaced, evolving to lower fluid to rock ratios as the instability of the mid-latitude ice resulted in sublimation and dessication of the regoltih, through to dry debris flows and infrequent reactivation of gullies by low volumes of liquids in the present climate. In the future, the spectral identification of water or concentrated salts in a recently reactivated gully might be attained, but there is also a major role to be played by simulating the flow of fluid in a martian regolith under current atmospheric conditions to attempt to recreate the fluvial structures seen on the surface [Conway et al., 2011b; Price et al., 2011]. The astrobiological distinction between a source fluid that originates subsurface in an environment that promotes long term water stability, versus seasonally available and transient near-surface liquid, was discussed through the distinction between active and dormant life in 4.1.2.
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Table 5-2: Activity and duration of recent flow features. Flow type
Latitude
Season of activity Autumn, winter, early spring
Orientation preference*
Physical context
Deposit albedo
Temp. (K)
Reference
Dune gullies
40-60
Equator
Sand, ice
Dark
< 150
48-32
Spring, summer
Equator
Dark
> 250
Crater wall gullies
30-55
Spring, autumn, winter
Pole
Bright
No
DDSs** and VFFs**
75-85
Spring, summer
Steep rocky slopes Originated from rocky alcove, transports small boulders Sand, ices
[Diniega et al., 2010, 2011; Dundas et al., 2010, 2011a, 2011b; Hansen et al., 2011; Jouannic et al., 2010] [McEwen et al., 2011a; Ojha et al., 2011] [Dundas et al., 2010; Harrison et al., 2009; Malin et al., 2006]
RSLs/ TSLs**
Dark
< 200
*Grey
shading indicates values not found in the literature. slope lineae; dark dune spots; viscous flow features
** Recurring/transient
[Kereszturi et al., 2011; Mohlmann and Thomsen, 2011; Reiss et al., 2010]
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Figure 5-10: Fresh bright deposits in Gasa crater gullies at 129.4 ˚E, 35.7 ˚S (HiRISE image12). New flows were observed in several gully channels within this system in the autumn to spring [Dundas et al., 2011a].
5.3.4.b.
Dune flows
Dune gullies (or dune flow features; DFF’s) and recurring slope lineae (below) have become the focus of recent attention due to the ~ 30 cm resolution of the HiRISE camera providing a wealth of information on these narrow features. A significant attribute of both these features is that they are topographically isolated (hills, dunes, domes, etc.) and so would not be connected to any local groundwater system (as in the gully aquifer model). Dune gullies are a distinct class from the
12
HiRISE image ESP_012024_1440; metadata in Appendix A, Table A1; processing details in Appendix B.
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more commonly studied crater wall gullies discussed above, as many of them are on a smaller scale with thin channels (< 10 m across) although they are morphologically similar, showing alcoves, channels and aprons [Reiss et al., 2007] (Figure 5-11). They also display longer channels on average than crater wall gullies [Balme et al., 2006]. Recent surveys of dune gullies have found that many systems are seasonally active in the winter (where CO2 ice covers the surface) with their channels and aprons growing and extending, suggesting that they are linked to the seasonal accumulation and sublimation of carbon dioxide frost rather than to water [Diniega et al., 2010, 2011; Dundas et al., 2010, 2011a, 2011b; Jouannic et al., 2010]. This is consistent with the activity seen on northern polar dunes with gullies forming during the sublimation of the seasonal CO2 polar cap [Hansen et al., 2011]. Other dune gully systems given the term ‘viscous flow features’ (Table 5-2) show activity during the spring and summer when temperatures exceed the frost point of CO2 and there is no CO2 ice within the soil. These features are more likely a combination of debris flows lubricated by thin films of liquid or water, meltwater from H2O frost, or brines with low eutectics (cryobrines) [Kereszturi et al., 2011; Mohlmann and Thomsen, 2011; Reiss et al., 2010].
5.3.4.c.
Recurring slope lineae
Recurring slope lineae (RSL’s, formerly TSL’s) are defined as narrow (few metres across), linear features distinguished by their dark albedo that contrasts with the brighter slopes [Ojha et al., 2011] (Figure 5-12). They have been observed on dunes and crater walls and derive the term ‘transient’ from their fading (brightening) in subsequent imagery within a martian year. These features may have a fluvial origin, as seasonal (active during summer), latitudinal (mid-southern latitudes, primarily equator facing slopes) and slope relationships (steep rocky slopes > 20˚) indicate their association with warmer temperatures. They also display a fluvial morphology, with digitate fingers and diversion around topographic obstacles [Ojha et al., 2011]. Shallow brines have been proposed as a plausible formation agent [McEwen et al., 2011a] with the liquid derived from pore spaces in the regolith (hence the association of the features with porous rocky slopes rather than sandy and dusty regions) or from atmospherically trapped water vapour. It remains difficult, however, to distinguish a dry debris flow from a slightly lubricated flow through modelling, as the morphology and scale produced
135
by both is comparable [Pelletier et al., 2008]. Furthermore, the absence of RSL’s identified in the northern hemisphere requires explanation, and neither salts not water have been spectrally identified in fresh RSL’s but work is being undertaken in that direction. Mapping of water vapour behaviour during summer over slopes that display RSL’s and slopes that do not within a similar latitude band, would help to constrain an atmospheric origin for the putative fluid. Results presented in Chapters 6-8 will be later utilised in Chapter 9 (9.2) to present a formation theory for these features and predict their spatial distribution.
Figure 5-11: Active dune gullies in the Kaiser dune field at 46.7 ˚S, 20.1 ˚E [Diniega et al., 2010] (HiRISE image13). The red rectangle on the right indicates the extent of the left image.
13
HiRISE image ESP_016907_1330; metadata in Appendix A, Table A1; processing details in Appendix B.
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Figure 5-12: Active recurring slope lineae observed in the 298 km Newton crater at 41.6 ˚S, 202.3 ˚E.
Image is a composite of HiRISE observations and 3D slope
modelling. Flows appear seasonally in spring and summer and are potentially evidence of shallow liquid brines [McEwen et al., 2011b]. The narrowest flows are ~ 1.4 m across. Image Credit: NASA/JPL-Caltech/Univ. of Arizona14.
14
http://www.nasa.gov/mission_pages/MRO/multimedia/pia14479.html
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5.3.5. Beneath the cryosphere The ground penetrating radar MARSIS has not detected a subsurface aquifer in the top ~ 600 m of martian regolith (its sensitivity limit for an aquifer [White and Stofan, 2010; White et al., 2009]). If aquifers exist on Mars, they would most likely occur at the base of the martian cryosphere – the permanently frozen region of the crust- many kilometres beneath the surface. The present-day thickness of the cryosphere is estimated to average ~ 7 km at high latitudes and ~ 2 km at low latitudes with a nominal water ice content of (5-14)×107 km3 [Clifford, 1987, 1993; Hanna and Phillips, 2005]. This model however may not be consistent with some observations – for example, the 250 km diameter Lyot crater with a depth range of 15-30 km would be expected to have excavated beneath the cryosphere however it presents no evidence of hydrologic activity [Russell and Head, 2002] other than the presence of gullies in the central crater peak [Hart et al., 2009; McEwen et al., 2010]. Incorporating brines can allow the base of the cryosphere to occur nearer to the surface, for example at high latitudes the base of the cryosphere could occur at 4.1 km depth for a brine with freezing temperature of -21 ˚C consistent with NaCl [Grimm et al., 2006]. As discussed in 5.1.2 and 5.3.4.a, a liquid aquifer beneath the cryosphere would only be stable with respect to vapourisation if the pore space above it is blocked by ice [Grimm and Painter, 2009].
5.4.
Summary
Under the current climate with its low atmospheric pressure and frequent subzero temperatures, it is difficult to envision the action of liquid water. Nevertheless, with the benefit of repeated satellite observation of surface environments and the increasing data from surface missions, a story of ongoing subtle liquid water activity is developing. Liquid water is a necessary ingredient for Earth life and hence any liquid water environments on Mars are of strong astrobiological significance. In this review it was shown that evidence from radar sounding of ice, hydrogen concentrations from neutron flux, and spectral observations of water absorption present a generally consistent story of subsurface water. The low latitudes are generally dry but enriched localised regions of subsurface water occur ~10’s of centimetres beneath desiccated material and may have remnant
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Literature Review & Background: Liquid Water on Mars
water ice from periods of higher obliquity (for example, within Arabia Terra). At mid- to high-latitudes water ice is stable within millimetres of the surface and may extend deep into the subsurface in many locations (such as radar sounding of eastern Hellas Planitia). Water frost is deposited on the surface or shadowed slopes during the wintertime at most latitudes. Water ice environments might experience seasonal melting (at the surface or at the depths of gully alcoves ~200 m beneath) and produce recent flows and new deposits in gully, dune gully, and RSL systems. Salts are globally abundant on the surface and presumably within the regolith, extending the temperature range and stability of potential liquid water environments.
In Chapter 6 the martian surface and subsurface P-T
environment will be modelled to identify locations that have conditions favourable for melting. The predictions of martian liquid water environments will then be compared to the locations of water ice and putative liquid water presented here and will be utilised to predict the occurrence of transient liquid water morphologies.
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6. PHASE DIAGRAM OF MARS, WATER & THE POTENTIAL MARTIAN BIOSPHERE
‘No one would have believed in the last years of the nineteenth century that this world was being watched keenly and closely by intelligences greater than man's and yet as mortal as his own; that as men busied themselves about their various concerns they were scrutinised and studied, perhaps almost as narrowly as a man with a microscope might scrutinise the transient creatures that swarm and multiply in a drop of water. With infinite complacency men went to and fro over this globe about their little affairs, serene in their assurance of their empire over matter. It is possible that the infusoria under the microscope do the same. Yet across the gulf of space, minds that are to our minds as ours are to those of the beasts that perish, intellects vast and cool and unsympathetic, regarded this earth with envious eyes, and slowly and surely drew their plans against us.’ -
H.G. Wells, ‘War of the Worlds’
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Phase Diagram of Mars, Water & the Potential Martian Biosphere
The search for life on Mars can be informed by knowledge of inhabited water on Earth. In Chapters 3 and 4 it was shown that liquid water can occur on Earth to a maximum depth of ~ 70 km from pressure-temperature constraints, but that only 19 % of the volume of the Earth within the 70 km thick shell of the hydrosphere is known to harbour life. It was shown that the strongest limiting factors for life within liquid water environments are temperatures above ~ 122 ˚C and low water activity < 0.6. Here a pressure-temperature model of Mars is developed, compared to the P-T diagram of martian water chemistries, and used to quantify the maximum possible extent of liquid water on Mars. The P-T conditions of Earth environments that are hospitable to life (quantified in Chapter 4) are then superimposed on Mars to identify potentially hospitable liquid water. The depths where these potentially hospitable liquid water environments can be encountered on Mars, and the likelihood of water with sufficient liquid activity for life in such environments, is then examined. Any liquid water on Mars that is in contact with the martian atmosphere will be transient.
The low partial pressure of water vapour below the triple point
pressure of 610 Pa results in rapid sublimation or vapourisation of exposed water, leaving the surface and shallow porous regions of the martian subsurface generally dry (5.1.2).
In comparison, water vapour is a larger constituent of Earth’s
atmosphere and the partial pressure is generally above the triple point (Table 6-1) so water is predominantly stable. Carbon dioxide composes 95 % of the martian atmosphere [Litvak et al., 2004], remains stable throughout a wide range of P-T conditions that occur on the surface [Leighton and Murray, 1966], and due to its low temperature liquid regime (Figure 1-3) can potentially occur as a fluid in some subsurface environments [Stewart and Nimmo, 2002].
Although there are
theoretical arguments that non-aqueous solvents can potentially support life [Badescu, 2011b], the focus in this thesis is on liquid water as it is the only solvent with proven astrobiological significance.
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Table 6-1: Average pressures and temperatures on Earth and Mars Environment Pressure Ptot (Pa)
Partial pressure
Temperature
of water Ppart (Pa)
(˚C)
Earth surface
105 (3×104 - 1.2×105) 104 – 102
15
Mars surface
6 x 102 (70 - 1.2×103) 2 – 10-2
-60
6.1.
Development of model
6.1.1. Average brine & largest liquid phase space From the discussion in Chapter 5, liquid water on Mars would be exposed to a number of different salts to those found in aqueous solutions on Earth (Table 5-1). The presence of these soluble compounds in the martian soil suggests that saline solutions may exist in the martian subsurface if liquid water is present [Clark and van Hart, 1981]. Additionally the colder martian climate, reaching lows of 140 K at the mean surface [Lodders and Fegley, 1998], provides a wide range of temperatures within the regime of subzero liquid brines and thin films. Using an approach analogous to that in Chapter 3, the pressure-temperature range of water chemistries that are most relevant to martian environments are modelled. Through synthesising the discussion in section 5.3.3 on the detection of salts on the martian surface and assuming that these salts pervade the crust, liquid water on Mars would likely be a brine containing 8–30 % salt by mass. The incorporated salts would include magnesium and calcium sulfate, sodium chloride, and magnesium and calcium carbonate [Boynton et al., 2009; Jagoutz, 2006; Marion et al., 2009b, 2010; Tosca et al., 2008].
Given the information on the eutectic
temperatures of these salts in Table 5-1, it was estimated that a representative martian brine would have a weakly depressed freezing point in the range 253-263 K, depending on the number of ions in the system [Jagoutz, 2006].
This
conservative estimate of an ‘average martian brine’ with a freezing point of 263 K has a corresponding boiling point elevation of 2.8 K from Equation 2-2. From shifting the phase boundary curves of pure water the triple point pressure for the martian brine is ~ 320 Pa (intersection of the vapourisation and melting curves).
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Phase Diagram of Mars, Water & the Potential Martian Biosphere
The liquid range for the most prevalent martian salts was not calculated individually as they all fall within the liquid range of the estimated ‘average martian brine’, and so are in agreement with the results presented below. The exceptions to this are sodium chloride and perchlorates, which have lower freezing points but have not been found to occur in high/eutectic concentrations within the soil. The full envelope of P-T conditions on Mars that allow water to remain a liquid corresponds to the ‘maximum liquid range’.
Thin liquid films can persist to
temperatures of as low as 173 K [Pearson and Derbyshire, 1974] and potentially lower (estimates of 143 K [Mazur, 1980]) and hence they are the extreme boundary of liquid water phase space. This liquid regime may occur in association with water ice in martian permafrost or coating grains in the soil during high humidity [Mohlmann, 2008]. Using a freezing point depression of -100 ˚C in Equation 2-1 gives a solute molarity of 53.8 moles which provides a boiling point elevation of 27.5 ˚C from Equation 2-2. The triple point occurs at a pressure of 0.004 Pa.
At its minimum temperature, the phase space represents the coldest
possible liquid water on Mars. The maximum liquid range incorporates not only thin films under the full range of martian surface conditions, but also potentially concentrated brines, such as a perchlorate brine with a freezing point depression of 70 K [Marion et al., 2010; Zorzano et al., 2009].
6.1.1.a.
Stability & subsurface closure
The closure depth of the martian regolith is very important for liquid water duration, as from the discussion in 5.1 any water environments that are exposed to the martian atmosphere will be unstable and their volume will vary strongly with the atmospheric partial pressure. A higher Ppart in the atmosphere relative to the soil pore space results in vapour diffusing into the soil to be absorbed as liquid films and condense as ice (when temperatures fall below the frost point). A lower Ppart will cause exposed ice to sublimate into the atmosphere [Sizemore and Mellon, 2006]. Ice within the pore space provides an impermeable layer and hence blocks the substrate below from having contact with the atmosphere [Schorghofer, 2010]. Any water between the surface and the ice table depth however will be unstable or metastable and will fluctuate diurnally or seasonally [Haberle et al., 2001]. The
143
depth to perennial ice varies between 0 cm at the poles (a small permanent water ice cap exists at both poles) to ~ 1 m at 50˚ and > 3 m or no stable ice-table at the equator [Chevrier et al., 2008; Mellon and Jakosky, 1995], generally decreasing with increasing latitude (5.2.3). More generally, when ice is not present to provide an impermeable layer, the stability of subsurface martian water will be determined by the depth of pore space closure. The permeability of the martian subsurface to air is not well constrained (further details in 6.2.2.a). Fine grained particulates such as dust and sand effectively fill the pore spaces in the regolith [Hudson and Aharonson, 2008]. Salt too can reduce pore space by cementing soil grains together [Jakosky and Christensen, 1986]. In the absence of ice, the closure depth of the martian subsurface will therefore vary strongly with substrate. Soil on Earth can be permeable to the atmosphere for tens of centimetres before a sealing agent (such as salt) or an impermeable strata is encountered [Heard et al., 1988]. On Mars the situation is likely similar. The effects of compaction cannot be guaranteed to isolate pore space from the surface until ~ 2-4 km depth [Clifford, 1993; Schumacher and Zegers, 2011; Zegers et al., 2010]. This provides an important cautionary caveat to the interpretation of the state of water in the martian subsurface from the phase model. Although the vertical axis in the P-T model of Mars derived below estimates the subsurface lithostatic pressure environment, this can only be equated to the pore space pressure environment when the pore space is isolated (see also discussion in 3.1.7.b). For example, a temperature of 293 K at a depth of 10 m falls within the liquid regime of an average martian brine, however without the presence of a sealing agent or impermeable strata the shallow depth indicates that the environment will be connected to the surface atmosphere, and hence any water will not be stable. The same temperature at a depth of 2 km may well indicate a potentially long-term liquid water environment. Even if in diffusive contact with the atmosphere however, the duration of any unstable liquid water will increase with burial depth due to the longer diffusion timescale for vapour [Mellon et al., 1997].
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Phase Diagram of Mars, Water & the Potential Martian Biosphere
6.1.2. Mars model To examine the full range of pressure and temperature environments on Mars the modelling was divided into five sections: atmosphere, crust and surface, mantle and core. Knowledge on the P-T conditions within the atmosphere and the surface is predominantly from direct observations, through infrared (thermal) and radio sensing from orbit, and measurements from the six successful surface probes. Hence the variability of pressures and temperatures in this region of Mars are well constrained.
Within the crust, mantle and core however, pressures and
temperatures are primarily derived from geophysical models and scaling relations to the Earth and Moon. The main data sources for models of the P-T conditions in the martian interior are: -
Mineralogies and densities of primitive martian material (SNC meteorites), and in-situ measurements on the surface, to constrain the mineralogy, state and temperature profile of the mantle and core [McSween et al., 2009]
-
Gravitational field measurements (geodesy), to constrain mass and moment of inertia [Rivoldini et al., 2011a, 2011b]
-
Gravity and topography, to constrain crust thickness and average density [Wieczorek and Zuber, 2004]
-
Magnetic field measurements, to constrain the state of the martian core dynamo [Lillis et al., 2008]
-
Thermal imaging to identify crustal hot spots and heat flow [Christensen et al., 2003]
It is beyond the scope of this work however to discuss how geothermal models are developed, and how the above datasets have evolved the understanding of Mars’ deep P-T environment. Instead this work is constrained to utilising the estimates of pressures and temperatures at depth, and their uncertainties, from the range of published models. The pressure and temperature conditions within the aforementioned regions of Mars (atmosphere, surface, crust, mantle and core) were modelled in two ways, analogously to the Earth model in Chapter 3. Firstly, representative conditions within each region were illustrated through an average geotherm– the most likely temperature conditions encountered at a given pressure or depth, due to the outwards flow of heat from Mars’s interior and the thermal conduction of the
145
dominant materials. Secondly, to represent the uncertainty in geothermal models and the variation in martian materials, the width in P-T space of each region encompasses the full variation in geothermal models and measured conditions. For example, the uncertainty in the boundary conditions in the central core, coremantle boundary, and crust-mantle boundary (Moho) are represented through vertical and horizontal error bars encompassing the range of estimates in the literature.
Within the crust and atmosphere however, the shading is more
representative of the range of actual conditions on present-day Mars as there is a more complete and accurate set of P-T measurements available.
6.1.2.a.
Pressure & depth
The pressure experienced in Mars’ interior is related to the subsurface depth through: Equation 6-1
ܲ ൌ ͲͲ ͵Ǥ͵ሺߩଵଵ ݖଵଵ ߩଵଶ ݖଵଶ ߩଶ ݖଶ ߩଷ ݄ݖଷ ሻ
This pressure is due the overburden of atmosphere (the 1st term) and the layers of crust (2nd term and 3rd term), mantle (4th term) and core (5th term). The mean density ɏ (kg/m3) within each layer will be given below. 600 Pa is the total pressure Ptot at zero elevation on Mars, z is the depth under the surface (m) and
3.73 m/s2 is the gravitational acceleration at the surface (g). The variation in gravitational acceleration was not included in this model (unlike the Earth model, 3.1.2), as due to the large uncertainties in estimates of Mars’ radial mass distribution [Folkner, 1997], estimates of the variation of g with depth are not available. . 6.1.3. Atmosphere Table 6-2 gives the conversion used from altitude above the martian surface to pressure, and the mean geotherm used to plot the range of temperatures experienced with altitude in the atmosphere.
To estimate the variation in
temperature at a given altitude within the atmosphere, the mean geotherm was
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Phase Diagram of Mars, Water & the Potential Martian Biosphere
translated to the maximum and minimum temperatures measured through radio occultation at the highest elevation surface – Olympus Mons15 [Hinson et al., 2004]. Table 6-2: Model parameters for Mars’ atmosphere. Parameter Mean lapse rate temperature (K/km) Lapse rate pressure (Pa/km) References
Value 1.1 8.6 [Barth, 1974; Conrath et al., 2000; Lodders and Fegley, 1998]
6.1.4. Surface & crust 6.1.4.a.
Boundary
The extreme environments of the surface of Mars are given in Table 6-3 and used to define the pressure-temperature boundary of the surface region.
These
environments have predominantly been directly sampled/measured (points A, B, C, HB, PL). Table 6-3 also includes point D which has been extrapolated to a lower pressure than the original estimate.
Due to entry decent and landing (EDL)
constraints on martian surface missions [Grant et al., 2011], there is a lack of surface temperature measurements below zero elevation. Hence the measurement made by Opportunity at 3 km below the zero-elevation surface [Smith et al., 2006] is used as the maximum surface temperature at zero elevation. With the exception of HB and PL, the points in Table 6-3 represent extreme P-T conditions, and hence all environments on the surface (a subset of the ‘crust’ region) must lie within the pressure-temperature polygon defined by this boundary. To define the crust boundary at higher pressures than the surface region, two approaches were used. Firstly, the letters E, F and G in Table 6-3 represent plausible estimates of some extreme P-T conditions from geophysical models. Secondly, the maximum (25 K/km) and minimum (4 K/km) plausible geothermal gradients in martian crustal rocks were included (6.1.4.c below).
For the hottest plausible subsurface
temperatures, the maximum geotherm was applied beneath the hottest temperature conditions at zero elevation. For the coldest plausible subsurface temperatures, the minimum geotherm under the coldest surface was used. This 15
http://nova.stanford.edu/projects/mod/
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provided a boundary for the temperatures that could be encountered at depth between 0-50 km (the outer region of the crust). It is assumed that all the points within this crust region exist on Mars (see also 3.1.4.a). To represent the mean conditions within the crust and hence the most likely P and T conditions that would be encountered at depth, the variation in crustal thickness, heat flow, and geothermal gradients will be examined. Table 6-3: Extreme temperatures and pressures on Mars’ surface and crust Key
Environment
A,B
Olympus Mons summit
C
Minimum surface temperature (South Pole ice cap) Maximum surface temperature (Opportunity landing site, 3 km altitude) Shallow, cold Mohorovicic discontinuity at base of crust Deep, cool Mohorovicic discontinuity
D*
E
F
Pressure (Pa) Depth (km) 25 -27 3.9×102 -5
Temp. Reference (K) 135-177 [Hinson et al., 2004] 140
[Carr, 2006; Kieffer, 1979]
3.9×102 0
303
[Carr, 2006; Smith et al., 2006]
3.8×108 35
350
[Wieczorek and Zuber, 2004; Zuber et al., 2000]
1.1×109 100
610
G
Deep, hot Mohorovicic discontinuity
1.2×109 100
HB
Lowest point on surface, maximum surface pressure (Hellas Basin, 29˚S) Northern arctic in late spring/ early summer (Phoenix Landing site)
(0.9-1.02)×103 8.2
PL
7.5×102 2
[McGovern, 2002; Sohl et al., 2005; Wieczorek and Zuber, 2004] 920 [McGovern, 2002; Sohl et al., 2005; Zuber et al., 2000] 150-203 [Hinson et al., 2004]
193
[Nelli et al., 2009]
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6.1.4.b.
Phase Diagram of Mars, Water & the Potential Martian Biosphere
Mean & variation
Present estimates of Mars’ crustal thickness are based on indirect geophysical studies involving local relationships between gravity and topography [Sohl et al., 2005] (Figure 6-1). Estimates of crustal thickness range from 35-100 km [Turcotte et al., 2002; Zuber, 2001] with a global average of 50 km [Neumann et al., 2004; Wieczorek and Zuber, 2004].
This range is considerably smaller than earlier
estimates placing the thickest regions of the crust at 200 km with a typical value of 100 km [Sohl and Spohn, 1997]. These earlier estimates predated the more recent revised topography and gravity results from Mars Global Surveyor [Zuber et al., 2000]. Lithostatic pressure estimates within the crust, due to mass overburden, assume an average crustal density within the range of 2700-3100 kg/m3. The former is representative of the andesitic-basaltic composition from Pathfinder measurements of soil free rocks [Wieczorek and Zuber, 2004]. The latter, from basaltic shergottites in SNC meteorites interpreted as martian crust samples [Britt and Consolmagno S.J., 2003]. Outside this range of mean crustal rock densities, there is a wide density variation associated with different substrates within the crust. For example, the polar ice caps are estimated to have a density within 10002000kg/m3 [Johnson, 2000], depending on the fraction of dust and sediments. Density also increases with load, and hence a higher density would be expected at zero elevation beneath the Tharsis volcanoes compared to the mean surface [McGovern, 2002]. In this model, a mean crustal density of 2900 kg/m3 is used for crust < 50km thick [Neumann et al., 2004; Turcotte et al., 2002; Zuber, 2001], consistent with current geophysical models [Cheung and King, 2011; Zharkov et al., 2009]. For crust 50-100 km thick a higher load density of 3100 kg/m3 is used [McGovern, 2002]. Hence from Equation 6-1, this results in a range of (0.4-1.3) x 109 Pa at the base of the crust between 35-100 km depth (Table 6-4). The temperature at the base of a mean-thickness martian crust (50 km) is modelled as 450 K (180 ˚C) [McKenzie et al., 2002; Verhoeven et al., 2005], consistent with a geotherm of ~ 5 K/km in rock (6.1.4.c). Some older estimates for the crust-mantle boundary temperature lie between 1300-1700 K [Lodders and Fegley, 1998; Sohl and Spohn, 1997] however these models utilised the older estimates of 200 km thick crust. The range of plausible crustal thicknesses results
149
in a range of Moho temperatures, modelled as 350-920 K (Table 6-3). These extreme temperatures for the base of the crust, from the mean surface (210 K) to crust between 35-100 km depth, result in average geothermal gradients across the crust between 4-7 K/km, consistent with estimates for present day Mars [McGovern, 2002]. The mean P-T conditions within the crust and the Moho are represented by an average geotherm. Error bars encompass the variation (due to the range of crustal thicknesses and heat flows) of typical P-T conditions at the base of the crust.
Figure 6-1: Mars Global Surveyor gravitational anomaly map [Smith et al., 1999]16. Gravitational anomalies are regions where the gravity and topography are not correlated, and can be used to infer subsurface densities and elastic thickness [McKenzie et al., 2002].
16
http://www.astro.lsa.umich.edu
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Phase Diagram of Mars, Water & the Potential Martian Biosphere
Table 6-4: Model parameters for Mars’ crust Parameter
Mean
Thickness (km)
50
Temperature (K)
Surface: 210 Base: 450 Surface: 600 Base: 5.4×108 35-100 majority in 35-62 average ~35 in NH; ~58 in SH; globally 50 2900 (0-50 km) 3100 (50-100 km) 4.8
Pressure (Pa) Depth (km)
Density (kg/m3) Geothermal gradient (K/km) References
6.1.4.c.
Range (incorporating measured variation and model uncertainty) 35-100 Surface: 140-300 Base: 350-920 Surface: 250-1200 Base: (0.4-1.3)×109
2700-3100 4-25
[Britt and Consolmagno S.J., 2003; Cheung and King, 2011; Hahn and McLennan, 2008; Hauck II and Phillips, 2002; Hoffman, 2000; Lodders and Fegley, 1998; McGovern, 2002; McKenzie et al., 2002; Neumann et al., 2004; Phillips et al., 2008; Ruiz et al., 2006, 2010; Sohl and Spohn, 1997; Sohl et al., 2005; Solomon and Head, 1990; Turcotte et al., 2002; Urquhart and Gulick, 2003; Verhoeven et al., 2005; Wieczorek and Zuber, 2004; Zharkov et al., 2009; Zuber, 2001; Zuber et al., 2000]
Geothermal gradients in the crust
Below the depth at which the temperature varies with the daily and seasonal cycles of solar heating at the surface (~ 10’s of metres [Jakosky, 1983], further details in Chapter 7), the subsurface temperature profile is dominated by the outwards geothermal heat flow from the core. As there has never been a direct estimate of the martian heat flow the temperature environment is not well constrained, however scaling of Earth’s values can be carried out under the assumption that smaller planetary masses have lower core and mantle temperatures and lower surface heat flow [Papuc and Davies, 2008]. As on Earth, the net heat flow on Mars is the summation of contributions from the dissipation of energy from the core and the decay of radiogenic elements in the crust and mantle [Ruiz et al., 2004].
Radioactive elements U, K and Th have been spectrally
identified on the martian surface, with their concentrations leading to an estimate
151
of the present day heat flow from radioactivity of ~ 6 mW/m2 [Hahn and McLennan, 2008].
Information on the distribution and stratification of these
elements within the crust is lacking. Geothermal heat flow estimates for Mars are predominantly derived from the elastic thickness of the lithosphere and hence the depth to the brittle-to-ductile transition (assuming a thermal conductivity for the crust). The thickness of the elastic lithosphere is defined by the depth at which ductile behaviour replaces brittle behaviour [Solomon and Head, 1990], which is related to the vertical thermal gradient within the crust and mantle [Brace and Kohlstedt, 1980].
It is determined by comparing the gravity and stresses
associated with large lithospheric loads [McGovern, 2002] however the derived heat flow estimates are valid only for when the structures formed. The youngest significant lithospheric structures, which provide the most accurate estimates of the present-day heat flow, are the polar caps [Ruiz et al., 2010, 2011]. Estimates for the present-day total surface heat flow on Mars are in the range 5-25 mW/m2 [Hauck II and Phillips, 2002; McGovern, 2002; Phillips et al., 2008; Ruiz et al., 2006, 2010; Solomon and Head, 1990; Urquhart and Gulick, 2003]. Heat flow can be converted to a geothermal gradient using thermal conductivity through Equation 3-2. Assuming a typical bulk thermal conductivity of martian crustal rocks of 2 W/m/K [Grott and Wieczorek, 2011] and a surface heat flow of 10 mW/m2, provides an average geothermal gradient of ~5 K/km. There are a number of caveats however in the application of Equation 3-2. Firstly, given that the distribution with depth of radiogenic elements within the martian crust is unknown and could account for > 50 % of the total heat flow [Ruiz et al., 2006], the reliable application of the heat flow equation is limited to the near surface. Secondly, thermal conductivity varies widely with substrate and temperature conditions [Piqueux and Christensen, 2011]. The P-T boundary values of the crust used in the Mars phase model result in an average geotherm of 5 K/km throughout the crust. Significant variation in subsurface geothermal gradients within the martian crust is expected, as observed on Earth, due to the variation in thermal conductivity of crustal materials. The thermal conductivity variation is most significant in the shallow subsurface < 4 km where the crust is more heterogeneous [Zuber et al.,
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Phase Diagram of Mars, Water & the Potential Martian Biosphere
2000]. For example, water and CO2 ice which compose the polar caps have a significantly lower thermal conductivity than rock [Clifford, 1987; James, 1968]. Taking a range of 0.001-4.5 W/m/K for martian materials from dust [Mellon and Phillips, 2001] to rock [Clauser and Huenges, 1995] (for reference, a value of 4 W/m/K is generally applied to the martian mantle [Breuer and Spohn, 2003]), and the variation in heat flow estimates of 5-25 mW/m2, gives a range of geothermal gradients from 1.1-2.5×104 K/km valid for the shallow subsurface. The extremely high geotherms correspond to layers of low conductivity materials (sand and dust) which only constitute a very thin layer on the martian surface compared to the bulk of the planet. For example, dust deposits can be 200 m thick [Christensen and Moore, 1992] but are typically < 5 m [Christensen, 1986a] and sand in martian dunes can form deposits tens of metres high [Bourke et al., 2006] (more discussion on shallow temperature gradients in Chapter 7). For reference, geotherms in crustal rocks on Earth vary predominantly between 5-70 K/km (3.1.4.c). Although the mean geotherm in the Mars P-T model is ~ 5 K/km, the full width of the shaded region is consistent with geotherms up to 25 K/km across 50 km of crust.
6.1.5. Mantle & core Mars’ mantle is modelled as extending from a depth of 35-100 km beneath the surface depending on the thickness of the crust, to 1650 km depth [Mocquet et al., 2011; Sanloup, 1999].
The lower boundary is defined by the transition
temperature of the FeO-MgO-SiO2 system [Bertka and Fei, 1997] and so is strongly dependent on mantle and core composition (with Fe content being the most variable between models). Variations in mineralogy estimates and independent geophysical constraints, such as tidal deformation, lead to an uncertainty in the depth of the base of the mantle within the range of 1300-1900 km [Folkner, 1997; Longhi et al., 1992; Sanloup, 1999; Sohl and Spohn, 1997; van Thienen et al., 2006; Yoder et al., 2003]. For the core-mantle boundary (CMB) a representative pressure of 2.1×1010 Pa is used [Sohl and Spohn, 1997], consistent with a depth of 1650 km, and the densities described in Table 6-5. Model-dependent temperature estimates for the CMB range between 1500-2000 K however a representative average temperature of 1700 K is utilised here as it is favoured by more recent ‘cool’ models [Fei and Bertka, 2005; Khan and Connolly, 2008; Verhoeven et al., 2005]. Given the modelled conditions at the Moho and CMB boundaries, the average
153
geotherm across the martian mantle is ~ 0.8 K/km (0.7 K/km in the Earth’s mantle, Table 3-5). Table 6-5: Model parameters for Mars’ mantle and core Mantle
Mean
Depth (km)
Outside: 50 Inside: 1650 Outside: 450 Inside: 1700 Outside: 5.4×108 Inside: 2.1×1010 3100 (50-100 km) 3500 (100-1650 km) 0.8
Temperature (K) Pressure (Pa) Density (kg/m3) Geothermal gradient (K/km) References
Core Temperature (K) Pressure (Pa) Depth (km) Density (kg/m3) Geothermal gradient (K/km) References
Range (incorporating measured variation and model uncertainty) Outside: 35-100 Inside: 1300-1900 Outside: 350-920 Inside: 1500-2000 Outside: (0.4-1.3)×109 Inside: (1.6-2.5)×1010
[Bertka and Fei, 1997; Fei and Bertka, 2005; Folkner, 1997; Khan and Connolly, 2008; Lodders and Fegley, 1998; Longhi et al., 1992; McGovern, 2002; McKenzie et al., 2002; Mocquet et al., 2011; Sanloup, 1999; Sohl and Spohn, 1997; van Thienen et al., 2006; Verhoeven et al., 2005; Wieczorek and Zuber, 2004; Yoder et al., 2003; Zuber et al., 2000] Mean Range Outside: 1700 Outside: 1500-2000 Inside: 2200 Inside: 1950-3000 Outside: 2.1×1010 Outside: (1.6-2.5)×1010 10 Inside: (3.8-5)×1010 Inside: 4.1×10 Outside: 1650 Outside: 1300-1900 Inside: 3390 7300 0.7 [Bertka and Fei, 1997; Dehant et al., 2011; Fei and Bertka, 2005; Folkner, 1997; Gudkova, 2004; Khan and Connolly, 2008; Lodders and Fegley, 1998; Longhi et al., 1992; Sohl and Spohn, 1997; Stewart et al., 2007; Verhoeven et al., 2005]
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Phase Diagram of Mars, Water & the Potential Martian Biosphere
Mars’ core is modelled as extending from a depth of 1650 km to 3390 km [Dehant et al., 2011]. Core size and temperature profile are estimated from Mars’ polar moment of inertia [Folkner, 1997] which is measured from orbit. This parameter indicates the amount of central condensation in the planet and so can constrain the combination of core sizes and masses/densities that are possible. In the centre the core pressure is modelled as 4.1×1010 Pa [Sohl and Spohn, 1997]. Estimates for this pressure lie predominantly between (3.8-5)×1010 Pa [Gudkova, 2004; Sohl and Spohn, 1997] (Table 6-5), with a particular low pressure estimate of 2.4×1010 Pa [Fei and Bertka, 2005]. Hence, the average value used in the model is reasonably representative. The estimates vary widely as the poorly constrained composition of the inner core leads to significantly different model geotherms [Gudkova, 2004]. Core composition is estimated predominantly from the isotopic composition of SNC meteorites [Lodders and Fegley, 1998].
Central core temperatures are
modelled to lie between 1950-3000 K however recent modelling favours the cooler end of this range around 2200 K [Bertka and Fei, 1997; Fei and Bertka, 2005; Khan and Connolly, 2008]. Hence 2200 K is used as the representative central core temperature in this model. This is significantly cooler than the central Earth temperature of 6100 K, highlighting the significant cooling of Mars’ core and consequently the lower heat flow across the planet’s surface. The phase of the present-day martian core is dependent on its exact mineralogy. A liquid core is consistent with its moment-of-inertia [Barlow, 2008] and data from mineral transition experiments for the estimated core composition indicate that the martian interior is liquid as its geotherm is at a higher temperature than the melting isotherm [Bertka and Fei, 1997; Stewart et al., 2007]. However, Mars exhibits only a very weak magnetic field indicating that it does not have an active dynamo and hence its inner core must have either completely solidified or be liquid in very slow motion [Longhi et al., 1992]. With the chosen boundary values in the Mars P-T model, the average geotherm across the core is 0.7 K/km, comparable to the models for Earth’s core geotherm (0.85 K/km, Table 3-5). The temperatures and pressures along the average model geotherm leads to an entirely liquid core for typical model compositions [Fei and Bertka, 2005]. As the mineralogy and hence the conditions within the martian mantle and core are poorly constrained it is expected that these regions of the model will be likely
155
to need revision as more data becomes available and uncertainties are reduced [Dehant et al., 2011; Mocquet et al., 2011]. The representative values used for the central core and the transition between core and mantle, as well as the range of uncertainties in these values discussed above, are given in Table 6-5.
6.1.6. Interpreting the Mars phase diagram 6.1.6.a.
Transient environments
In the Earth pressure-temperature model (3.1.6), some environments were classified as ‘transient’ to reflect the dynamic nature of the Earth.
This
encompassed short term environments at crustal pressures but mantle temperatures. Mars has been a dynamic planet and has experienced volcanism in its geologically recent past. Many of the martian lava flows have absolute ages in the range 200-500 Ma [Berman and Hartmann, 2002; Hiesinger et al., 2007], with some examples potentially as young as 10-50 Ma [Hartmann and Berman, 2000; Hartmann, 1999]. Given that shield volcanoes can lie dormant for 100 Ma, it is therefore plausible that Mars is still geologically active [Mitchell and Wilson, 2003] but has not displayed evidence during the period of human observation. It is also likely that Mars once supported hydrothermal environments much like the Earth. The presence of minerals such as montmorillonite [Clark et al., 2007] and kaolinite [Mustard et al., 2008] on the surface suggests the interaction of heated water with surface rocks in the past. The existence of current hydrothermal environments cannot be definitively ruled out [Schulze-Makuch et al., 2007].
A speculative
indicator of current high temperature water or magma environments on Mars is the global observation of methane at an average of 10 parts per billion (ppb) [Formisano et al., 2004], despite a photochemical lifetime of methane in the martian atmosphere of only 340 years [Krasnopolsky et al., 2004]. A number of Earth-based observations also reported localised enhancements reaching ~ 30 ppb of methane [Mumma et al., 2009] (Figure 6-2). If these observations of a modern abundance of methane on Mars are real (the measurements have been contested and follow-up studies are required [Encrenaz, 2008; Zahnle et al., 2011]), then one possible explanation is outgassing from magmatic or hydrothermal systems [Atreya et al., 2007], for example, the methane generated from high temperature aqueous reactions with iron oxide [Krasnopolsky, 2006]. However, despite the
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Phase Diagram of Mars, Water & the Potential Martian Biosphere
evidence for shallow, hot environments in the martian past, there is no definitive evidence that these environments currently exist. Thus there are no transient environments to include in the P-T model for Mars.
Figure 6-2: Global atmospheric methane concentration on Mars. Three regions of enhanced methane have been observed in Terra Sabae, Nilli Fossae and Syrtis Major [Mumma et al., 2009].
6.1.6.b.
The meaning of pressure
The dual meanings of the vertical pressure axis in the Earth and Mars P-T models were discussed in 3.1.7.b. For Mars’ environments, the relevant pressure is the overburden pressure Ptot of either the atmosphere (on the surface) or rock (beneath the surface, Equation 6-1). For interpreting the phase of water present in a martian environment, the partial pressure Ppart is relevant (5.1). Difficulty in interpretation arises when the partial pressure is not known as it is dependent on whether the pore space is connected to the atmosphere or to a source of water (such as ice) within the subsurface (6.1.1.a).
A number of situations can be
examined. The discussion in Chapter 3 demonstrated that it is reasonable to use lithostatic pressure for all environments below the closure depth (including subsurface liquid water in pore space) as the density of rock in Mars’ crust and upper mantle is less than three times the density of water [Lodders and Fegley,
157
1998].
This density assumption therefore introduces a minimal error to the
pressure value compared to the inherent uncertainty in the values used to construct the model. This reasoning is supported by models of subsurface fluid pressures on Mars which indicate that putative fluids may be at near-hydrostatic pressures down to depths of ~ 10 km [Hanna and Phillips, 2005]. Hence, beneath the perennial ice table depth a mid- to high latitudes (6.1.1.a), or the unknown closure depth of pore space at low latitudes, the pressure values on the vertical axis can be used to approximate both the total pressure in the subsurface pore space and the partial pressure of water vapour if a source of water is present.
6.1.7. The P-T space of Earth’s life The phase model of Earth’s biosphere was constructed in Chapter 4 by identifying the extreme pressures and temperatures of environments known to be inhabited by life (listed in Table 4-1 and Table 4-2). In this chapter the phase space of life is superimposed onto the Mars P-T model, to investigate whether there are environments on Mars that could potentially be habitable for life assuming sources of chemical energy, nutrients and liquid water were available. The ‘potential martian biosphere’ (PMB) is here defined as all environments on Mars where the pressure and temperature are within both the growth and metabolism thresholds of Earth organisms, and within the liquid region of the phase space of an average martian brine. As the focus here is to obtain an upper bound to the extent of potentially habitable martian environments, it is necessary to include those P-T conditions that are plausibly suitable for life but have simply not been sampled on Earth. Hence high P conditions above the maximum pressure of ~108 Pa sampled for life but within the 253-395 K (-20 ˚C to 122 ˚C) range for active life are included within the PMB. The validity of this extrapolation was discussed in section 4.1.1. Also included within the PMB are environments which do not exist on Earth, but may be habitable given the current understanding of the P-T thresholds of life and the results of Chapter 4. These consist of low pressure environments that occur within the martian subsurface but are at pressures below the ~ 104 Pa of the highest altitude active life on Earth. These environments arise on Mars due to the lower surface pressure (~ 6/1000 of Earth at zero elevation/sea level) and the shallower pressure gradient in the subsurface (gravity is ~ 1/3 that of the Earth). Ideally environments that have low T conditions below the minimum temperature
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Phase Diagram of Mars, Water & the Potential Martian Biosphere
of 253 K for life, but that are within liquid water with an activity above 0.6 and hence may support active life would also be included.
The occurrence of
environments that support dormant life on Earth will also be investigated on Mars.
6.2.
Results: the pressure-temperature representation of Mars &
liquid water The phase diagram of water chemistries relevant to Mars is shown in Figure 6-3. The Mars P-T model shown in Figure 6-4 presents the regions of phase space that correspond to martian environments, both measured and estimated from geophysical models. The phase diagram of martian water is superimposed, to identify where in phase space there is both Mars and the potential for liquid water. This is useful for investigating the extent of habitable environments on Mars. Figure 6-4 also reveals regions where there is Mars but no liquid water, and where there could be liquid water but there is no Mars.
Figure 6-3: Phase diagram of water relevant to martian environments. Triple point of the largest water phase space is 173 K and 0.004 Pa. The triple point of an estimate average martian brine is 263 K and 320 Pa.
Figure 6-4
159
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Phase Diagram of Mars, Water & the Potential Martian Biosphere
On the previous page: Figure 6-4: Mars and liquid water. Subsurface seasonal temperature variations in rock at absolute latitude 30˚ are shown in grey (see Chapter 7). The horizontal thin line at 600 Pa is the average atmospheric pressure of Mars. The black cross on that line marks the average surface temperature of 210 K. Letters indicate the extreme PT values listed in Table 6-3. The black diamonds give the mean modelled conditions at the crust-mantle boundary, core-mantle boundary, and the centre of the core. The red asterisk represents the hottest and deepest possible liquid water on Mars at 647 K and 3.2×109 Pa, corresponding to a depth of ~ 260 km. The red triangles indicate the intersection of the minimum, average and maximum martian geotherm with 395 K, the current maximum temperature for life.
6.2.1. Implications of high solute & thin film liquids The solid curves in Figure 6-3 mark the estimate of the maximum liquid range on Mars, corresponding to thin liquid films in ice and permafrost and potentially concentrated briny inclusions [Mohlmann and Thomsen, 2011; Mohlmann, 2010] (6.1.1). The dotted curve marks the conservative estimate of a martian brine composed of < 10 mol/kg of salt [Marion et al., 2009b] or < 30 % salts by mass [Clark and van Hart, 1981]. On Mars, the presence of P-T conditions within the stable or metastable region of liquid water does not imply that water is present within that environment. From the discussion in section 5.2, the presence of widespread high latitude volatiles in the shallow subsurface can be inferred, as can the localized presence of volatiles in the low latitudes, however there is little knowledge of the spatial distribution of subsurface water on Mars or its depth extent. Hence, although physical and thermodynamic arguments on the maximum and plausible extent and volume of the martian hydrosphere will be presented, this work is limited by the aforementioned caveats. As the salts present on Mars produce only a moderate freezing point depression when in solution (Table 5-1, an exception to this is perchlorate salts which were found to be important at the Phoenix lander site [Hecht et al., 2009] but whose global abundance is unknown [Marion et al., 2009a]). Martian salts therefore will likely produce only a minor increase in the temperature space of liquid water,
161
compared to pure water. From Figure 6-4 the representative martian brine can be metastable as a liquid at the surface, when Ptot > 320 Pa and T > 263 K, only in a small fraction of environments with the warmest surface temperatures. These will be in regions of the highest solar insolation, in equatorial latitudes. High humidity or a source of water ice is also a necessity for the actual presence of liquid water in these environments to be realised.
The P-T conditions for liquid brine
metastability continue into the subsurface beneath these surfaces if there is a high subsurface geothermal gradient. A geotherm of 20-25 K/km beneath a surface within the liquid brine regime allows for subsurface liquid brines- indicated by the intersection of the high temperature crust region with the liquid regime. The conditions under which such a high geothermal gradient can occur within the top 10’s of metres of the martian regolith will be examined in Chapter 7. From the average geotherm of 5 K/km in Figure 6-4, it is clear that typical subsurface temperatures are too cool to allow for shallow liquid brine. The mean geotherm does not reach the melting isotherm of 263 K until a depth of 10.3 km (1.12×108 Pa), indicating that this depth is likely to be more representative of the upper boundary of a potential global hydrosphere, with the occasional regional exception of a shallower liquid brine. Beneath this depth liquid water can be perennially stable. This estimate provides a deeper depth for the upper boundary of a liquid aquifer than previous models, which used higher heat flows of 30-50 mW/m2 to give estimates within the range of 2-7 km [Clifford, 1987, 1993; Hanna and Phillips, 2005] (5.3.5). It is worth mentioning however that there is currently no observational evidence for deep liquid water on Mars and the 10.3 km depth of the 263 K melting isotherm is far beneath the deepest known deposits of water ice from radar: ~ 4 km beneath the local surface at the poles [Holt et al., 2008; Phillips et al., 2008; Plaut et al., 2009] (5.2.3) and 1 km deep in the Cerberus region [Picardi et al., 2005]. Hence these locations will support deep subsurface liquid water only if ice extends to the melting depth. The martian subsurface is heterogenous, and hence there are likely to be locations where the melting depth is encountered either shallower or deeper than the 10.3 km estimate for the hydrosphere depth.
This variation in the depth to the
hydrosphere is due to variations in the crustal geothermal gradient, in response to
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Phase Diagram of Mars, Water & the Potential Martian Biosphere
variations in thermal conductivity and heat flow (Equation 3-2). As discussed previously (6.1.4.c), the thermal conductivity of martian materials ranges broadly from 0.001-4.5 W/m/K [Mellon and Phillips, 2001], however estimates of the unmeasured modern day martian heat flow vary over a much smaller range of 5-25 mW/m2 [McGovern, 2002; Ruiz et al., 2010].
Hence the variation in thermal
conductivities within the crust has a much larger effect on the deep geothermal gradient than the heat flow. To estimate the plausible range in depths of the upper boundary of the hydrosphere, the extreme values of heat flow and thermal conductivity can be used. From Figure 6-4, melting P-T conditions – indicated by the intersection of the liquid phase space with the crust region – can occur much deeper than 10.3 km (the depth at which the mean geotherm intersects the liquid regime). Taking the lowest heat flow estimate of 5 mW/m2 and the highest thermal conductivity of 4.5 W/m/K corresponding to plutonic rocks [Clauser and Huenges, 1995], results in a geothermal gradient of 1.1 K/km and the 263 K isotherm occurring at a depth of 48.2 km beneath the surface (consistent with Figure 6-4). Estimating the shallowest plausible melting depth is more complex. Figure 6-4 suggests melting P-T conditions can be encountered at or just below the martian surface if geothermal heat dominates and the surface layer has sufficiently low thermal conductivity [Heldmann and Mellon, 2004; Mellon and Phillips, 2001]. This would occur in a surface layer composed of small grain size materials [Presley and Christensen, 1997], such as dust and sand.
However, shallow subsurface
temperatures within the ~ 10’s of metres depth of the annual thermal wave are determined more strongly by the seasonal variations in solar heating, and the parameters which affect the absorption of that heat (such as albedo), rather than the deeper geothermal heat flow [Pollack and Huang, 2000].
The specific
conditions that allow temperatures of 263 K at the martian surface or in the shallow subsurface in different materials, will be examined in Chapter 7.
In
summary, from current geophysical estimates and the known properties of materials in the martian crust, the depth to the potential global hydrosphere of ଷǤଽ Mars is estimated as 10.3 േଵǤଷ km.
163
From Figure 6-4, thin films occupy a much larger region of P-T space than a typical brine. They can occur throughout the full range of elevations on the martian surface for some duration during the year (as the triple point occurs above the highest elevation surface), and are likely to be the dominant state of shallow liquid water on Mars [Mohlmann, 2009b]. Furthermore unlike a weak brine, they can be stable when exposed to the atmosphere, as their triple point pressure of 0.004 Pa is less than the Ppart of water vapour (0.01-2 Pa [Melchiorri et al., 2007, 2009; Smith et al., 2009b]). The ideal conditions for formation of liquid films are high humidity and the presence of salts on the surface to absorb water vapour and form a thin liquid covering through deliquescence [Searls et al., 2010]. Unfrozen thin films in the martian soil are predicted to increase in volume with depth throughout the extent of the regolith that is connected to the atmosphere [Grimm et al., 2006; Mohlmann, 2005]. Figure 6-4 reveals that liquid water films (or concentrated brines) can also occur in the martian atmosphere. Water ice in clouds has been observed at ~20 km above the surface [Nelli et al., 2010; Whiteway et al., 2008] and dust particles have been observed up to ~25 km [Tamppari et al., 2010]. Thus there is the potential for thin films of liquid water on particulate grains in the martian atmosphere analogous to that observed in Earth’s troposphere [Amato et al., 2007].
6.2.2. Deepest liquid water The deepest and hottest liquid water on Mars as predicted by the P-T model is shown by the red asterisk in Figure 6-4 (also Table 6-6). It corresponds to the intersection of the maximum pressure of Mars’ phase space with the critical point of the estimated martian brine. This environment occurs at 647 K (374 ˚C) and 3.2×109 Pa, corresponding to a depth of ~ 260 km and hence lying within the mantle (Equation 6-1). As this is the maximum extent of liquid water, it represents the lower boundary of any potential biosphere on present-day Mars. Liquid water is stable at slightly higher pressures than this limit, but given the estimated subsurface geothermal gradient throughout the crust and mantle of Mars, environments at higher pressures below the critical point of water do not exist. The maximum extent of the martian hydrosphere corresponds to at most 21 % of the volume of Mars with the capacity to support liquid water (Table 6-7). The
164
Phase Diagram of Mars, Water & the Potential Martian Biosphere
corresponding value for the volume of Earth within the hydrosphere is ~ 3 % (3.2.2). Table 6-6: Pressure and depths of features of interest. Description Critical point of pure liquid water along mean geotherm
Hottest possible liquid water (ignoring pore space)
Hottest pore space closure estimate, along mean geotherm
Hottest pore space closure estimate, at maximum depth
Pressure Temp. Pressure-depth (Pa) (K) constants 9 1.4 x 10 647 h11 = 5x104 ρ11 = 2895 h12 = 5x104 ρ12 = 3100 h2 = 2.2 x 104 ρ2 = 3500 3.2 x 109 647 h11 = 5x104 ρ11 = 2895 h12 = 5x104 ρ12 = 3100 h2 = 1.6 x 105 ρ2 = 3500 9 1.2 x 10 623 h11 = 5x104 ρ11 = 2895 h12 = 5x104 ρ12 = 3100 h2 = 6.2 x103 ρ2 = 3500 9 2.8 x 10 623 h11 = 5x104 ρ11 = 2895 h12 = 5x104 ρ12 = 3100 h2 = 1.3 x 105 ρ2 = 3500 *Relevant
6.2.2.a.
Depth (km) 121.5
259.4
106.2
228.8
to Equation 6-1
Pore space & fluid access
Despite subsurface P-T conditions that allow water to be liquid, the existence of liquid water and life will depend on the availability of pore space. Subsurface pore space is described in detail in 3.2.2.a but its relevance to Mars will be mentioned here. The maximum porosity at a given depth decreases with increasing depth as pore space becomes more restricted due to overburden pressure [Athy, 1930]. On Mars, a minimum threshold for pore space would be achieved at a greater depth than on Earth given the lower overburden pressure, as Mars has ~ 1/3 of Earth’s surface gravity and the subsurface pressure gradient scales accordingly. Pore space deforms at high temperatures of ~623-673 K [Hyndman and Shearer, 1989],
165
and hence rocks that have experienced burial depths of >100 km on Mars (from the depth of 623 K along the mean geotherm) would be expected to have no pore space. Overburden pressure in combination with the low yield strength of rocks due to high temperature may effectively close pore space before these temperatures are reached. Table 6-7: Volume estimates for the potential martian hydrosphere Description
Thickness of spherical shell (km) Mars 3390 Maximum extent of the hydrosphere 259.4 Maximum extent of the hydrosphere 228.8 considering pore space restrictions
Fraction of Mars’ volume % 100 21.2 18.9
Fraction of Mars’ volume within hydrosphere % ---100 89.0
Life may have a minimum pore space requirement [Drancourt et al., 2003; Knoll, 2003] (2.8.1), however it is likely that permeability - the connectedness of pore space which governs the flow of fluid- is also fundamental [Fisher, 1998]. Depth is the strongest control on subsurface permeability [Brace, 1980] with the strongest decrease occurring at the brittle-ductile transition (BDT) [Ingebritsen and Manning, 1999] due to the decreased fracturing and porosity of the rocks at this depth. On Mars, the BDT for crustal rocks is estimated to occur between 30-60 km depth [Montesi and Zuber, 2003; Ruiz et al., 2006]. Below the BDT, permeability is not a strong function of depth.
A medium is considered impermeable if its
permeability is < 10-19 m2 [Wyckoff et al., 1933]. Studies into the absence of flow morphologies at the 3 km deep Lyot crater constrained fluid flow in the deep rocks to moderate, with permeabilities in the range of 10-13-10-16 m2 [Russell and Head, 2002]. For comparison, a typical permeability in basaltic oceanic crust is 10-16 m2 [Manning and Bird, 1995].
Independent estimates suggest that the martian
subsurface permeability is minimal below 26 km [Clifford and Parker, 2001] (based on gravitational scaling of the fracture closure pressures on Earth) and hence the downwards percolation of groundwater could be restricted beneath this depth. Deeper pore space beneath the melting isotherm could be saturated with liquid however if the pore space was originally in contact with a source of ice.
166
6.3.
Phase Diagram of Mars, Water & the Potential Martian Biosphere
Results: the potential martian biosphere
The phase space of the potential martian biosphere is shown in Figure 6-5, superimposed on the phase space of martian environments and liquid water. All currently known life on Earth inhabits the pressures and temperatures shaded in light green, with the cross-hatching reflecting dormant life below 253 K (-20 ˚C). Given the demonstration in 4.2 of the observed high pressure limit to life on Earth being a selection effect, and the observed low pressure limit of life on Earth being due to limitations in nutrients and water activity (4.2.3), the dark green shading indicates extrapolated pressure conditions outside the known environments of Earth life, shown in Figure 6-6, but within life’s temperature thresholds. This shading reflects additional plausibly habitable environments on Mars. Thus Figure 6-5 shows that there are extensive environments in the martian crust that have PT conditions potentially hospitable to life. Some Earth environments do not exist in the martian subsurface (Figure 6-6), namely hot crustal environments described by the numbers 2-4. These environments are absent from Mars due to its smaller geothermal heat flow and lack of present day geologic activity. However, the lower mean geothermal gradient in the martian crust (~ 5 K/km compared to ~ 25 ˚C/km on Earth) provides deep, cool environments down to ~10’s of kilometres that are not found on Earth (Figure 6-6). The maximum depth for liquid water on Mars is ~ 260 km. Assuming that liquid water extends throughout environments that have P-T conditions within the liquid regime, and assuming there is a source of nutrients and energy available to subsurface life, then high temperature and low pore space are the key limiting factors for Mars’ potential deep biosphere. The currently held high temperature limit for life is 395 K (122 ˚C) [Takai et al., 2008] and hence this temperature determines the lower limit of the potential biosphere on Mars. From Table 6-8, the intersection of this isotherm with the mean geotherm of the Mars P-T model occurs at a depth of ~ 37 km and is within the phase space of average salinity liquid water. This provides an estimate of the global mean depth to which life could potentially extend into the martian subsurface based on current geophysical models, if liquid water is available.
167
Similarly to the discussion of the depth to the potential martian hydrosphere (6.2.1), the heterogeneity of the martian subsurface results in a range of plausible depths for the 395 K isotherm. From the phase space model in Figure 6-5 the vertical width of the crust region around 395 K indicates that these temperatures can occur between 3 km and 80 km depth in the martian subsurface. This broad range is due to the variations in thermal conductivity and heat flow that are incorporated into the phase space model. To examine the validity of the estimated depth extent of the biosphere, the values of thermal conductivity and heat flow that would result in maximum depths greater than 80 km, or less than 3 km, can be examined. Such conditions would therefore result in estimates that are outside the model. For the 395 K isotherm to be encountered deeper than 80 km below the subsurface a shallow geothermal gradient is required, arising from a low heat flow and high thermal conductivity (from Equation 3-2). The lower bound to the martian heat flow estimates is 5 mW/m2 and the upper bound to the thermal conductivity of crustal rocks is 4.5 W/m/K (see section 6.1.4.c). Using these values provides a geothermal gradient of 1.1 K/km which results in the 395 K isotherm being encountered at 168 km depth. This is clearly outside of the martian phase space model in Figure 6-5. More generally, depth values greater than 80 km would occur for geothermal gradients < 2.3 K/km. This is unlikely to occur, as the minimum geothermal gradients – which would be expected in regions of extensive thickened crust in the southern hemisphere [Lewis and Simons, 2011; Zhong, 2011] - are ~ 2.6 K/km (using values from [Hahn et al., 2011]). Geothermal gradients < 2.3 K/km could potentially arise from an extremely low heat flow < 5 W/m/K and crustal thermal conductivity > 2.2 W/m/K, or extremely high thermal conductivity > 4.5 W/m/K and heat flow < 10.4 mW/m2. For the 395 K isotherm to be encountered shallower than 3 km beneath the martian surface, an extremely high geothermal gradient of > 61.6 K/km would be required. Such a steep thermal gradient could only occur in regions of low thermal conductivity or high heat flow. The former condition will be examined first. Taking the minimum estimate for the modern martian heat flow of 5 mW/m2 would require the thermal conductivity to be < 0.08 W/m/K for the geothermal
168
Phase Diagram of Mars, Water & the Potential Martian Biosphere
gradient to be > 61.6 K/km. Thermal conductivity values less than this are not unreasonable at the martian surface, as particles of fine dust and sand < 500 µm across [Presley and Christensen, 1997] will have thermal conductivities in the range of ~ 0.001-0.08 W/m/K [Mellon and Phillips, 2001]. Hence in regions of Mars where the surface is dominated by small grain sizes (examined in Chapter 8) throughout a significant extent of the subsurface, the 395 K isotherm could potentially at depths shallower than 3 km, which is outside of the 3-80 km range of the phase-space model.
The second situation mentioned above is a steep
geothermal gradient due to extremely high heat flow. Taking the upper thermal conductivity estimate of 4.5 W/m/K for rock (6.1.4.c), a geotherm > 61.6 K/km would necessitate a heat flow > 277.2 mW/m2. This is clearly unreasonable for modern Mars (for reference, the highest geothermal heat flows on Earth are in the region of 150 mW/m2, Figure 3-2). The extent of the biosphere is not only determined by the temperature constraint of 395 K. It is also sensitive to variations in pore space and closure depth of pore space, which affect the rate of pressure increase with depth and the available space for life. As discussed in 6.1.1.a, it is possible that in regions of depleted subsurface ice, such as at the equator, the subsurface may not be isolated from the atmosphere until a depth of ~ 4 km. In such locations, the rate of pressure increase with depth will be minimal across the top 4 km of the subsurface and, more significantly, subsurface liquid brines would be exposed to the atmosphere and hence metastable (5.1.2).
This would therefore imply that any estimates for the
biosphere occurring within the top 4 km depth would be less likely, as thin films condensed from atmospheric water vapour would be the main source of liquid water for life. As discussed in the previous paragraphs, for temperatures of 395 K – indicating the maximum depth of the biosphere - to be encountered within 4 km depth, small grain-size materials must exist as a thick layer on the martian surface. In such regions however it is unlikely that the subsurface would remain permeable to the atmosphere throughout the top 4 km due to the likelihood of the pore space being filled by the small grains, and the salt cementation of small grains that is frequently observed on Mars [Hudson and Aharonson, 2008; Jakosky and Christensen, 1986;
169
Schorghofer, 2010]. Hence the permeability of the subsurface pore space to the atmosphere is unlikely to provide a significant limit on the estimated depth extent of the biosphere. The closure of pore space could potentially be a limiting factor for life however, if life has a minimum pore space requirement (2.8). From the temperature deformation of rock discussed in 6.2.2.a, pore space on Mars is estimated to be closed at depths > 100 km (Table 6-6). This is outside of the range of depth estimates for the biosphere as temperatures at >100 km depth would most likely exceed 395 K and hence be restrictive for life. In summary, considering the thermal conductivity, heat flow, and pore space variations of the martian subsurface, the extent of habitable temperature conditions for life and hence a potential biosphere on present-day Mars is estimated to be 37േସଷ ଷସ km. Values outside this model are unlikely, but could
potentially occur in regions dominated by an extensive, thick layer of small grain size materials or regions with very low modern heat flow and high thermal
conductivity rock. As the radiogenic heat production and temperatures of the deep martian interior become better constrained, or the surface heat flow is directly measured [Grott et al., 2011], then the range in uncertainly of the 395 K isotherm will decrease. Table 6-8: Volume estimates for the potential martian biosphere Description
Temp. Mean pressure (K) (Pa) [range]
Current max. for life Proposed maximum for life Proposed maximum for life
395
*Using
423
523
4.0×108 [3.2×107 – 8.9×108] 5.0×108 [4.5×107 - 109] 8.9×108 [8×107- 2×109]
Mean depth (km)* [range] 37.0 [3.0-80.2] 46.2 [4.2-89.7]
% Volume Mars**
% Volume Mars within hydrosphere**
3.2
15.2
4.0
19.0
80.2 [7.4176.2]
6.9
32.6
Equation 6-1. Assuming spherical symmetry, rm = 3390, the hydrosphere extends to a maximum depth of ~ 260 km below the mean surface, and using Equation 4-1 and Equation 4-2. **
170
Phase Diagram of Mars, Water & the Potential Martian Biosphere
On the next page: Figure 6-5: The potential martian biosphere. Superposition of active and dormant life environments from the Earth on phase space of Mars. Numbers 1-8 circumscribe the region occupied by active life (light green shaded). Some hot low pressure terrestrial environments do not exist in the martian subsurface, such as numbers 2-4. Numbers 9-14 represent the region occupied by dormant life (light green hashed). Two regions have been labelled as extrapolated life as they overlap the martian phase-space and liquid water and are within the temperature range of life, but are at either higher or lower pressure than known inhabited environments on Earth. The orange line and yellow line indicate the change that would occur in this diagram if life could extend to 423 K and 523 K respectively. The red asterisk represents the hottest and deepest possible liquid water on Mars.
Figure 6-5
171
172
Phase Diagram of Mars, Water & the Potential Martian Biosphere
Figure 6-6: Comparison between the phase space of Earth and Mars. Shallow mantle intrusions on Earth provide shallow, hot environments not found on Mars due to its low present-day geological activity. The low geothermal gradient in the martian crust (modelled average of ~5 K/km compared to ~25 K/km on Earth), and its variation (shown by the width of the phase model at a given depth), provides deep, cool environments not found on Earth. Some conditions for surface life on Earth are reached in the shallow subsurface of Mars. The currently believed upper temperature limit of life on Earth may be the result of the limited exploration of high temperature environments in hydrothermal vents and the deep crust (4.2.4). It is plausible that in the future the temperature limit of life will be found to be higher, as microbiologist have suggested that the real limit may approach 423 K (150 ˚C) [Segerer et al., 1993] or 52 3K (250 ˚C) [Baross and Deming, 1983; Deming and Baross, 1993; Straube et al., 1990]. The latter is less likely due to possible protein and amino acid denaturation [Holden, 2008; White, 1984]. These two scenarios, summarised in Table 6-8, would result in the lower limit of the biosphere extending to 46.2േସଶǤଽ ସଶ km (mean and range;
orange line in Figure 6-5) and 80.2േଽ ଶǤ଼ km (yellow line in Figure 6-5) respectively.
The estimated average martian brine can remain a liquid over both these depth
173
ranges and so neither of these temperature limits can be discredited from P-T constraints. Both, however, would result in a significant extension of the depth of the potential biosphere of Mars. Subsurface pore space and the closure of fluid fractures in the martian subsurface are not well constrained, but could limit the availability of space and fluid access for subsurface life anywhere between 30 km and 230 km depth (the former from pressure closure of fractures, that latter from temperature deformation). Therefore it is plausible that these factors will not restrict many environments at temperatures below 395 K given the depth of the average geotherm, but they may restrict any life at higher temperatures. The maximum possible depth of liquid water of ~ 260 km is equivalent to a spherical shell of 21 % of the volume of Mars. By comparing the mean extent of the potential martian biosphere to the maximum extent of the martian hydrosphere, the volume of Mars with potentially inhabitable liquid water can be estimated. Considering that the average depth of 395 K (122 ˚C) is ~ 37 km, the potential biosphere corresponds to ~ 15.2 % of the volume of Mars with liquid water. In other words, at least ~84.8 % of the volume of Mars where liquid water could exist cannot support life, given the best geothermal models of the martian interior to date. This estimate is clearly poorly constrained- there is significant uncertainty in the temperature conditions in the deep martian crust (see Table 6-5) and there is currently no evidence on reservoirs of deep liquid water in the crust although they have been predicted beneath the polar deposits of water ice [Clifford, 1987]. Although the volumes discussed above are estimates, they provide a first quantification of the global potential of Mars to support subsurface life (new results to be published by [Méndez, 2012] will make an interesting comparison). Furthermore, the estimated fraction of hospitable liquid water on Mars is interesting with regards to drawing a comparison to Earth. In 4.2.5 the known extent of Earth’s biosphere was found to occupy 19 % of the volume of the Earth containing water, with 81 % of the volume of Earth where liquid water is stable (and most likely exists) not known to harbour life. The volume included in the average extent of the average martian biosphere is ~3.2 % compared to 0.7 % for life on Earth. Hence given the current understating of representative conditions within the martian crust, Mars may have the potential to support a substantial
174
Phase Diagram of Mars, Water & the Potential Martian Biosphere
subsurface biosphere with the proportion of Mars with habitable conditions for subsurface life being five times larger than that of Earth.
6.4.
Summary
It was shown that that there are many P-T environments in the martian subsurface that could be hospitable to terrestrial life. Liquid water on Mars can extent to a maximum depth of ~ 260 km given the current geophysical models of the P-T conditions within the martian interior, however pore space constraints and fracture closure could restrict the maximum depth of subsurface liquid percolation to anywhere between 30-230 km. The potential biosphere of Mars extends from the surface to a depth of ~ 37േସଷ ଷସ km. This depth is determined by the mean depth
of the 395 K (122 ˚C) isotherm which marks the hottest known temperature
environment of life on Earth, with the range due to variations in heat flow estimates and thermal conductivities of martian materials. Hence the extent of the PMB corresponds to a shell of ~ 3.2 % of the volume of the planet. The upper boundary of the biosphere occurs where temperatures exceed 253 K within the PT regime of liquid water as either brines or thin films. Although the conditions on the martian surface only allow for transient liquid brine, stable on diurnal and seasonal timescales, it is shown that there are environments in the surface and shallow subsurface that are potentially hospitable to life. These occur in lowlatitude locations with high thermal inertia and low albedo, where high humidity, exposed salts, and sheltered environments promote the deposition of frost, seasonal melting, and liquid duration. Deeper liquid brine environments beneath the average depth of the estimated 263 K melting temperature of a plausible martian sulphate+chloride brine at 10.3 km could potentially support life, however there is no current evidence for water ice extending to this depth or the presence of deep liquid water. Considering the maximum possible extent of liquid water on Mars, the ~ 37 km shell that may support Earth-like life on Mars corresponds to ~ 15.2 % of the volume of Mars that may have liquid water hospitable for life. At least ~ 84.8 % of the volume of Mars that may have liquid water is inhospitable to life.
Many shallow and deep subsurface environments on Mars should be
hospitable for psychrophilic, halophilic and anaerobic thermophillic life, provided liquid water and a source of chemical nutrients and energy are available.
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7. BRIDGING CHAPTER: THE INFLUENCE OF SURFACE MATERIALS ON THE SUBSURFACE ENVIRONMENT
‘...take the tip of the pencil and magnify it. One reaches the point where a stunning realization strikes home: The pencil tip is not solid; it is composed of atoms which whirl and revolve like a trillion demon planets. What seems solid to us is actually only a loose net held together by gravity. Viewed at their actual size, the distances between these atoms might become league, gulfs, aeons. The atoms themselves are composed of nuclei and revolving protons and electrons. One may step down further to subatomic particles. And then to what? Tachyons? Nothing? Of course not. Everything in the universe denies nothing; to suggest an ending is the one absurdity.’ -
Stephen King, The Gunslinger
176
Bridging Chapter: The Influence of Surface Materials on the Subsurface Environment
This chapter provides a bridge between the two complementary approaches applied in this thesis. The approach in Chapters 3, 4 and 6 provided a global assessment of the depth to and extent of potentially hospitable conditions within the martian subsurface. Predominantly, hospitable environments were found to occur in the deep crust, more than 10 km below the surface. The model also identified a region of possible shallow hospitable environments within the top 10’s of metres of the subsurface, occurring where both (i) high geothermal gradients between 20-25 K/km and (ii) sufficiently warm surface conditions are present. The warm surface temperatures at low latitudes satisfy the second condition, and can coincide with possible locations of subsurface water ice, such as within Arabia Terra [Atkins and Barlow, 2011] (5.2.3). This chapter will examine where on Mars these potential shallow hospitable conditions occur beneath the surface, and more specifically, how subsurface geothermal gradients much larger than the modelled mean of 5 K/km would arise. Unlike the deep subsurface, the shallow temperature profile within the top 10’s of metres is dominated by the downwards conduction of solar heat through the regolith on diurnal (10’s of centimetres) and seasonal (10’s of metres) timescales.
Obliquity variations penetrate much deeper into the
subsurface - 100’s of metres- but are poorly constrained and not considered here. The thermophysical properties of surface materials (thermal conductivity, density, heat capacity, and albedo) govern how heat is absorbed and distributed through their bulk, and hence determine the subsurface temperature and water environment. Therefore, to analyse the habitability of the shallow subsurface within ~ 20 m depth, a different approach is required and is applied in Chapter 8. In this chapter the factors that determine subsurface temperatures are discussed, and in particular, how they influence the martian subsurface environment.
177
7.1.
Background
7.1.1. Subsurface temperature profiles The temperature profile within the shallow regolith is determined primarily by solar insolation rather than deep geothermal heat flow [Pollack and Huang, 2000]. The range of temperatures experienced within a material at the surface of a planet can be understood through a heat-conduction equation [Turcotte and Shubert, 1982]:
Equation 7-1
డ்ሺ௭ǡ௧ሻ డ௧
ൌܦ
డమ ்ሺ௭ǡ௧ሻ డ௭ మ
where T is the soil temperature (K), at time t (s), depth z (m), and D is the thermal diffusivity (m2/s) given by:
Equation 7-2
ܦൌ ఘ
where ρ is the density (kg/m3), C is the heat capacity (J/kg/K) and k is the thermal conductivity (W/m/K) [Luo et al., 1992]. The relevant heat source here is solar, and it is assumed that there is no heat source in the soil. This assumption is reasonable over a small depth range as there is unlikely to be significant radiogenic heat within the thin ~ 10’s of metres surface layer to which this model will be applied [Vilà et al., 2010]. From Equation 7-1, the downward conduction of thermal waves is dependent solely on the thermal diffusivity of the medium, the time scales of interest, and a necessary boundary condition to solve the differential equation at the surface.
A more sophisticated model could incorporate soil
anisotropy through the horizontal heat transfer between adjacent materials in the soil [Huang et al., 2008], however this effect plays a much smaller role in determining the subsurface thermal environment than the variations in soil thermal conductivity [Luo et al., 1992].
An additional factor influencing soil
temperatures is the non-conductive heat transfer by water vapour, which can be significant in shallow, porous soils, however the effect of such can be accounted for by a broad range of thermal conductivities [Kane et al., 2001]. As thermal waves propagate through the subsurface, their amplitude diminishes exponentially with depth from the surface due to imperfect conduction and horizontal losses. The
178
Bridging Chapter: The Influence of Surface Materials on the Subsurface Environment
skin depth reflects the efficiency of thermal damping as it is the depth at which the temperature variations within the subsurface have decreased to 1/e of their surface value. It is given by:
Equation 7-3
ߜൌට
గ
ൌ
ூ
ఘ
ට
గ
where I is the thermal inertia discussed below (W/m/s½/K; 7.1.2) and P is the relevant timescale in seconds (e.g. daily, seasonally). Hence, materials such as rock and ice that have high diffusivity (Equation 7-2) will experience proportionally larger temperature fluctuations at a given depth compared to their surface range, than will low diffusivity insulating materials like dust and sand. A solution to the partial differential Equation 7-1 requires a surface boundary condition. Surface heat variations can be modelled as sinusoidal curves, given by: Equation 7-4
ܶ௦ ሺݐሻ ൌ ܶ οܶ݊݅ݏሺమഏ ሻ ು
where To is the average surface temperature and οܶ is the amplitude of variation
around To within the time scale P [Turcotte and Shubert, 1982]. Depth and time dependent subsurface temperatures are then given by the solution in the equation below [Hillel, 1982; Turcotte and Shubert, 1982].
Equation 7-5
ష
௭
ܶሺݖǡ ݐሻ ൌ ܶ οܶ݁ ഃ ݊݅ݏሺమഏ െ ఋሻ ು
Despite the simplifications in the model, this parameterisation provides a good fit with subsurface measurements of Earth soil temperatures [Wu and Nofziger, 1999]. Since the maximum and minimum of the sine term in Equation 7-5 are +1 and -1, the envelope of the maximum and minimum temperatures around the average temperature ܶ at depth is given by: Equation 7-6
ష
௫ ሺݖሻ ൌ ܶ േ οܶ݁ ഃ ܶ
179
Given the implicit dependence of subsurface temperatures in Equation 7-6 on the periodicity of surface temperature variation (through the skin depth; Equation 7-3), it is clear that the attenuation of thermal waves is frequency dependent, with longer period waves (such as seasonal) damping less with depth than shorter period waves (such as diurnal). If temperature changes occur within a time interval of P they will propagate ~ ξ ܲܦmetres into the subsurface. Diurnal and
seasonal variations typically cannot be detected at depths greater than metres or tens of metres, respectively on Mars, whereas long-term temperature trends (such as obliquity changes) can penetrate to hundreds of metres [Pollack and Huang, 2000].
Below this depth the geothermal gradient (6.1.4.c) determines the
temperature profile. The application of Equation 7-6 to model martian subsurface soil temperatures assumes that the temperature variations of the surface are known, which is generally only the case in locations that have been directly sampled by landing missions [Smith et al., 2006]. Numerous authors have modelled the diurnal and seasonal temperature profiles of the martian subsurface (Figure 7-1) [Mellon and Jakosky, 1993; Mellon et al., 2004b; Savijärvi and Kauhanen, 2008; Savijärvi, 1995; Schorghofer, 2005]. These models (see the ‘lookup’ model described in 7.1.2) used measured properties of the surface to derive the surface energy balance and used the temperature of the surface materials from thermal emission, to model how the surface heat is propagated to the subsurface (analogously to Equation 7-5). Predictions from these models are in excellent agreement with orbital estimates and surface measurements of ice table depth [Mellon et al., 2009a] (5.2.3). If the insolation environment at a location on the surface of Mars can be completely described (with the solar inclination, surface latitude and longitude, time of day, surface slope, and other parameters that determine the energy received by the surface being known [Thunholm, 1990]), then the temperature of the surface depends on the albedo and thermal inertia of the surface material [Mellon and Jakosky, 1993]. Thermal inertia and albedo are physical properties that provide partially independent constraints on the nature of surface materials, and their combined use for mapping geologic surface features on Mars has been applied extensively [Palluconi and Kieffer, 1981]. Their values can be determined by remote sensing of planetary surfaces at thermal and visual wavelengths. The
180
Bridging Chapter: The Influence of Surface Materials on the Subsurface Environment
thermal inertia and albedo of the martian surface and their relevance to the subsurface thermal environment is described below.
Figure 7-1: Example subsurface temperature profiles in dust at 59 ˚N [Mellon and Jakosky, 1993] (Fig. 6). The vertical dotted line marks the mean annual temperature, while the horizontal line marks ~ 4 annual skin depths
7.1.2. Thermal inertia The thermal inertia of a material is a measure of its ability to conduct and store heat [Mellon and Jakosky, 1993]. For a pure substance it is defined as:
Equation 7-7:
ܫൌ ඥ݇ߩܥ
where I is the thermal inertia in tiu, which is equivalent to J/m2/K/s½ in SI units [Putzig, 2006]. The product ρC is the volumetric heat capacity (k, ρ and C defined after Equation 7-2). From Equation 7-7, thermal insulators (materials with low thermal conductivity) will generally have low thermal inertia, so they rapidly heat and cool at their surface due to their poor ability to distribute heat through conduction. The thermal inertia values referred to in this study were derived from the Thermal Emission Spectrometer (TES) onboard Mars Global Surveyor (unless otherwise stated).
181
Figure 7-2: Histograms of the global albedo (top) and nighttime thermal inertia (bottom) from TES. The histogram of Mars’ global albedo contains three strong peaks – at 0.15, 0.23 and 0.27 (binwidth of 0.01)- and a small peak at 0.35. The last is caused by ice and the second last is caused by dust. The remainder are caused by mixtures of sand, rocks and duricrust with dust and ice. Thermal inertia shows two strong peaks – at 55 tiu and at 225 tiu (binwidth of 5). The former is caused by dust on Mars; the latter includes contributions from a range of materials. Thermal inertia values of planetary surfaces are given by a complex combination of particle size, bedrock outcrop and abundance [Putzig and Mellon, 2007a], and degree of cementation [Mellon et al., 2000]. For Mars, the values (Figure 7-2) are obtained by finding a best fit to the nighttime infrared brightness temperatures,
182
Bridging Chapter: The Influence of Surface Materials on the Subsurface Environment
within the wavelength range: 5.5 - 100 µm, using a subsurface heat conduction model with appropriate surface boundary conditions:
Equation 7-8:
ௌ
ோమ
ሺͳ െ ܣሻ ܿݏሺ݅ሻ ܨூோ ܮ
డ డ௧
ܫට
గ డ் ȁ ᇲ ೌ డ ᇲ ୀ
ൌ ߝߪܶௌସ
where the first term represents the solar flux absorbed at the surface (S = solar flux at the Earth’s solar distance; R = orbital radius of Mars; A = albedo; i = solar incidence angle); the second term represents the thermal radiation from the atmosphere received at the surface; the third term represents seasonal CO2 condensation (L = latent heat of sublimation of CO2; m = mass of CO2 frost); the fourth term represents subsurface heat conduction (I = thermal inertia; Pday = diurnal period; T = temperature; Z’ = depth in metres below surface, normalized to the thermal skin depth); the right hand side represents the energy radiated to the atmosphere (ε = emissivity of the surface; ߪ = Stefan Boltzmann constant; TS =
surface temperature) [Mellon et al., 2000; Putzig et al., 2005]. This model is
referred to as a ‘lookup’ model as the aforementioned parameters are either measured (such as the albedo A, and the infrared brightness temperature of the surface
డ் ȁ ᇲ ) డ ᇲ ୀ
or independently modelled (such as the radius R), with the
exception of the thermal inertia. The thermal inertia is determined by finding the best fit solution to Equation 7-8. Thermal inertia (Equation 7-7) is most sensitive to variations in thermal conductivity, which on Mars is strongly related to both grain size and the degree of cementation of materials. Density and heat capacity of materials on Mars generally vary by a factor of 2-3, however thermal conductivity varies by over 3 orders of magnitude [Neugebauer et al., 1971]. The dominance of thermal conductivity leads to general trends in thermal inertia governed by grain size: fine-grained, loosely packed materials have lower thermal conductivity and hence lower thermal inertia; while larger particles, such as rocks and ices, have higher values of both. This relationship with grain size is due to larger particles typically sharing a larger surface area with their neighbours and hence having a higher bulk thermal conductivity (valid on Mars for grains smaller than 1 mm [Jakosky, 1986]). Cementation of grains increases their effective thermal conductivity as heat flows
183
across the contact surface area. This property has a significant effect on the orbital thermal inertia, with a factor of ~ 8 increase in thermal conductivity resulting from ~ 1 vol% of cementation/bonding [Piqueux and Christensen, 2009] (7.1.5.c). Materials with low thermal inertia respond quickly to temperature changes and their surface temperatures will closely match the phase of diurnal temperature variations [Jakosky et al., 2005]. High thermal inertia materials are slower to respond to temperature changes so their surface temperature profile lags behind the diurnal variations. Hence at a given time of day, the materials on the surface can display vastly different temperatures: as much as 30 K between dust (24 tiu) and coarse pebbles (800 tiu) [Mellon et al., 2000]. Furthermore, the higher the thermal inertia of the material, the smaller the amplitude of its surface diurnal temperature variation, as the material is effectively dispersing the incident heat throughout its thermal bulk. The thermal inertia of the surface is also subject to large temporal variation. Seasonal cycling of surface ice and dust produces a > 50 tiu change in thermal inertia at most locations, and hence global maps represent only the median of the surface [Putzig and Mellon, 2007a]. Furthermore some materials, such as basalt, display a strong temperature dependence on their heat capacity with an increase of 75 % over the range of annual surface temperatures on Mars [Piqueux and Christensen, 2011]. Such materials therefore contribute strongly to the seasonal variation in surface thermal inertia. The variation of materials on the martian surface is at a finer scale than the spatial resolution of orbital spectrometers. The footprint of TES used to measure both visual reflectance (for albedo) and thermal emission (for thermal inertia) is ~ 3 km/pixel [Christensen et al., 1998]. This heterogeneity of the surface at the spatial scale of spacecraft observations complicates the interpretation of apparent thermal inertia both horizontally and vertically, and results in the apparent thermal inertia for a mixture differing fundamentally from that of either pure component if their contrast in thermal inertia is large [Putzig and Mellon, 2007b]. For example, a decrease in the sand:rock ratio from 50:50 to 30:70 within a pixel increases the maximum seasonal thermal inertia of the mixture from ~800 tiu to ~2000 tiu [Putzig and Mellon, 2007b]. The thermal inertia derived for each pixel is
184
Bridging Chapter: The Influence of Surface Materials on the Subsurface Environment
considered to be a linear weighted average of the true thermal inertia of the varying materials within the satellite’s field of view. The vertical resolution of the orbital thermal inertia measurements by TES are determined by the diurnal skin depth (Ɂ) of the material (Equation 7-3). Thus vertical layering (such as dust over rock, or desiccated material above the subsurface ice table) within several multiples of Ɂ also contributes to the orbital thermal inertia. For example, a 630
µm coating of dust over rock can decrease the bulk thermal inertia from ~ 700 to ~ 600 tiu [Mellon et al., 2008b]. Methods can be developed to utilise this effect to search for subsurface ice [Kuzmin et al., 2009]. Although vertical and horizontal
sub-pixel structure are the strongest caveats on the interpretation of orbital thermal inertia, other factors such as slope and atmospheric pressure (due to gas conduction of heat in porous soil [Fergason et al., 2006]) also play a minor role. Currently, it is not possible to simplify this complexity without multi-temporal observations of each surface pixel at different times of day, to more accurately fit the temperature profile of the surface and hence decipher component surface materials [Fergason et al., 2006]. Such a method which can currently only be employed at landing sites. The uncertainties associated with the derived orbital thermal inertia map employed in this study are discussed in Chapter 8 (8.2.1). Despite these caveats, the orbital thermal inertia values in Figure 7-2 show good agreement with surface values from landers (see 7.1.4).
7.1.3. Albedo Albedo is the fraction of incident visible/near-infrared (NIR) solar radiation between 0.3 – 2.9 µm that is reflected by the surface [Christensen et al., 2001a]. Bright materials have higher values of albedo and darker materials lower. For example, dark soils on Earth have albedo ~ 0.05 compared to ~ 0.75 for fresh snow [Ahrens, 2007]. Materials with higher albedo values have a lower maximum temperature than dark materials, with the difference being > 30 K over the range of martian materials [Mellon et al., 2000].
Albedo varies strongly with the
atmospheric redistribution of dust and seasonal condensation of CO2 and H2O [Byrne et al., 2008]. For example, the global dust storm in 2001 [Smith et al., 2002] brightened surface albedo values by as much as 0.1 [Putzig and Mellon, 2007a] and likely induced large perturbations in day- and nighttime surface temperatures [Wilson et al., 2007]. Bright regions on Mars are indicative of fine-grained surface
185
dust or ice [Bandfield and Smith, 2003; Rogers et al., 2007], whereas dark regions correspond to mixtures of sand, rocks, or duricrust (cemented sand sized grains) with smaller proportions of dust and ices (Figure 7-2). Low albedo regions have been found to have strongly homogeneous mineral compositions [Rogers et al., 2007], consisting of a mixture of minimally weathered basalt (termed surface type 1 [Bandfield et al., 2000]), altered basalt (surface type 2 [Karunatillake et al., 2006]) and hematite [Christensen et al., 2001b]. The movement of dust and ice leads to a strong seasonal variation in global albedo, both at latitudes above ±60˚, due to the seasonal polar caps and winter deposition of surface frost [Calvin and Titus, 2008; Kelly et al., 2006], and globally due to the redistribution of dust which brightens surfaces with albedo < 0.2 and darkens surfaces > 0.3 [James et al., 2007]. Albedo and thermal inertia are only partially independent as both are related to particle size. As discussed above, thermal inertia is related to thermal conductivity by a power law, which in turn is strongly related to particle size [Jakosky, 1986]. Albedo is also correlated with particle size, with surfaces composed of larger particles typically being darker due to lower specular reflection [Shkuratov et al., 1999]. Hence albedo and thermal conductivity of surface materials are generally negatively correlated. Furthermore, the albedo and thermal conductivity of a given soil vary inversely with soil moisture, with increased water within a soil lowering the albedo but raising the conductivity [Wang et al., 2005]. Therefore, on Mars, high thermal inertia materials (such as rocks) predominantly have lower albedos than small grained, low thermal inertia materials (such as dust, sand). A scatterplot of global thermal inertia and albedo values on Mars (Figure 7-3) reveals the complex relationship between these variables, showing a weak negative correlation between albedo and thermal inertia for values > 0.3 and < 250 tiu respectively (grain size < 400 µm). A third physical parameter useful for the characterisation of martian surface materials is their emissivity, e, defined as the ratio between the thermal radiation emitted by a substance and that emitted by a perfect black body at the same temperature. Based on this definition, the emissivity values of materials distributed on the surfaces of planets are < 1. Emissivity and visual albedo are
186
Bridging Chapter: The Influence of Surface Materials on the Subsurface Environment
negatively correlated for most materials, with the exception of ice, water and their mixtures. On Earth it has been shown that albedo and emissivity show a stronger (negative) correlation to each other than either does to the surface soil type [Zhou et al., 2003]. Similarly on Mars, emissivity is not useful for distinguishing amongst low albedo surface materials such as sand, duricrust and rocks, however it is useful in distinguishing bright dust and ice [Christensen, 1982].
Figure 7-3: Scatterplot of 2.592×107 pixels in the global map of median thermal inertia and albedo derived from thermal and visual reflectance data. The discrete values of thermal inertia below 50 tiu are an artefact of the ‘lookup’ model used to derive thermal inertia (Equation 7-8).
7.1.4. Groundtruthing Vertical layering of grain size can result in discrepancy between orbital and ground thermal inertia at each of the martian landing sites. For example, the layer of dark sand covering bright finer sand discovered beneath the top ~ 5-10 mm layer by Spirit at Home Plate [Arvidson et al., 2010], and beneath the top 5 mm surface layer at Columbia Hills [Wang et al., 2008] both lie within the depth sensitivity of TES and hence decreased the bulk thermal inertia to less than that of the surface layer.
187
Given that the global thermal inertia and albedo values (shown in Figure 7-3) are an average over ~ 9 km2 of surface area, their values are not expected to equal ground
observations.
Comparisons
between
remote
thermal
inertia
measurements and those determined at three of the six landing sites demonstrate that the unconsolidated soil component (fines) and the degree of grain cementation (duricrust) dominate the bulk thermal inertia determined from orbit (Table 7-1). Although rocks (> 2.5 cm diameter) have thermal inertia values an order of magnitude larger than fines, their surface coverage at each of the landing sites was < 20 % [Golombek et al., 2008b] and the thermal inertia derived from orbit was consistent with a combination of duricrust and fine-coarse soil. Orbital thermal inertia values consistent with rock are not therefore expected to be common. Hence, although the degree of layering and surface heterogeneity can be partially constrained by the combination of thermal inertia and albedo, further independent datasets such as visual imagery and shallow radar are needed to more conclusively decipher the fractions of different surface components. Nevertheless, orbital values of thermal inertia and albedo are generally consistent with the fine-component of the surface as measured by the landers (Table 7-1). Table 7-1: Comparison between ground and orbital thermal inertia and albedo Landing site
Ground* Thermal Albedo inertia
Orbit Thermal Albedo inertia 350-600 0.13-0.2 200 0.2
Curiosity** Phoenix
85-490
Opportunity
100-150
0.12-0.29
200
0.15
Spirit
175-400
0.2-0.3
290
0.19
283 234 386
0.22 0.24 0.19
Viking 1 Viking 2 Pathfinder
Reference
[Pelkey et al., 2004] [Mellon et al., 2009a; Putzig and Mellon, 2007a; Zent et al., 2010] [Bell et al., 2004a; Fergason et al., 2006; Golombek et al., 2005; Herkenhoff et al., 2008] [Bell et al., 2004b; Christensen et al., 2004a; Fergason et al., 2006; Golombek et al., 2008b] [Golombek et al., 2008b] [Golombek et al., 2008b] [Golombek et al., 2008b]
*Missing values occur where lander did not have the capability to measure surface reflectance. ** This mission is scheduled to land in August 2012
188
Bridging Chapter: The Influence of Surface Materials on the Subsurface Environment
7.1.5. Key martian surface components When examined together, martian albedo and thermal inertia values provide information on the distribution of dust, rocks and ice on the surface. These materials have extreme values in both parameters and hence are easily distinguished when the surface is cleanly exposed (Figure 7-2). Intermediate values of albedo and thermal inertia are more difficult to interpret, as they can represent either pure components of sand or duricrust, or sub-pixel mixing in varying proportions of end-member materials. The key component materials on the martian surface, distinguished by grain size, are well known from visual imagery and surface missions and are discussed below. Surface mineralogy is not considered in this work.
7.1.5.a.
Dust
The atmosphere of Mars is never dust free and the surface, globally, is dominated by fine particulates [Bandfield et al., 2000]. The global histogram of martian thermal inertia has a strong peak at ~ 55 tiu with a corresponding bright albedo peak at 0.27 (Figure 7-2) which is a signature of martian dust. A combination of infrared spectral analysis and studies of analogue materials at martian surface pressures reveals that dust on Mars typically has diameters less than ~ 40 µm [Christensen, 1986a] with a typical thermal inertia of 65 tiu [Jakosky, 1986], but ranging up to 120 tiu [Presley and Christensen, 1997] (Table 7-2). Albedo is strongly correlated with the degree of dust coverage [Ruff and Christensen, 2002] as martian dust is bright, with albedo > 0.27 [Kieffer et al., 1973], due primarily to its high red reflectance (Figure 7-4). The degree of martian dust coverage on the surface, defined by the dust cover index (DCI), can be determined independently by the average thermal emissivity at wavenumbers 1350-1400 cm-1 [Ruff and Christensen, 2002], with a low DCI indicating a high dust coverage.
189
Table 7-2: Model thermal inertia values of dust to sand sized particles. For range see Table 7-3. Particle size (µm)* 10 50 100 500 1000
Thermal conductivity (W/m/K)* 0.01 0.015 0.03 0.08 0.1
Density × specific heat** (J/m3/K) 106 106 106 106 106
Thermal inertia (tiu)*** 100 122 173 283 316
Approximate diurnal skin depth (cm)**** 1.8 2.2 3.1 5.0 5.6
* Values
taken from [Presley and Christensen, 1997] Product of density and specific heat on the martian surface is assumed to be 106 J/m3/K [Neugebauer et al., 1971] *** Calculated from Equation 7-7 **** Calculated from Equation 7-3 **
Figure 7-4: Histogram of pixels within the thermal inertia range of 5-100 tiu and their corresponding albedo values. From the location of the albedo modes, these pixels are consistent with pure dust > 2 cm thick (the diurnal skin depth), or a mixture of sand with varying proportions of dust.
7.1.5.b.
Sand
Larger, coarse, un-bonded particulates ranging from 0.1-4 mm dominate the low albedo regions of Mars [Christensen and Moore, 1992]. These intermediate size particles, representing sand and granules [Presley and Christensen, 1997], encompass a wide range of values of albedo and thermal inertia due to the relationship of both parameters to effective particle size [Mellon et al., 2000]. For sand-sized particles, thermal conductivity increases linearly with particle size
190
Bridging Chapter: The Influence of Surface Materials on the Subsurface Environment
[Jakosky, 1986] and the thermal inertia generally increases as the square root (Equation 7-7). Assuming martian sand particulate densities to be > 700 kg/m 3 [Presley and Christensen, 1997], thermal inertia values can be estimated (Table 7-3). For example, large sand grains of ~ 2 mm diameter have a thermal inertia value of ~ 400 tiu [Edgett and Christensen, 1994]. This represents the upper threshold for martian aeolian dunes, as they are generally composed of sand with grain sizes between 400-600 µm diameter [Edgett and Christensen, 1991], and hence regions covered in sand have thermal inertias ~150 -400 tiu. From Figure 7-5, the range of 150-400 tiu comprises materials with large variability in albedo, with two modes at 0.15 and 0.23, likely to be characteristic of pure sand (albedo 0.15) and sand beneath a layer of bright dust (albedo 0.26).
Figure 7-5: Histogram of pixels within the thermal inertia range of 150-400 tiu and their corresponding albedo values. These values are consistent with a layer pure sand > 4 cm thick, or a mixture of dust, sand, pebbles, and duricrust.
7.1.5.c.
Duricrust
In the martian environment, sand-sized grains can become cemented together by salt (sulphates and chlorides [Clark et al., 1982]), adsorbed water and ice [Arvidson et al., 2009]. Duricrust has been observed at all of the Mars landing sites [Mellon et al., 2008a] and may comprise a large fraction of the surface [Palluconi and Kieffer, 1981]. It is a weakly cohesive blocky soil [Christensen and Moore, 1992] composed of fine sand-sized 100-300 µm grains cemented together [Greeley et al., 2006; Jakosky and Christensen, 1986] with a moderate value of thermal inertia of ~ 430
191
tiu as measured on the surface [Christensen et al., 2004a], and ~ 350-450 tiu as measured from orbit [Moore and Jakosky, 1989; Squyres et al., 2004a]. Values up to 600 tiu may also be consistent with duricrust, as numerical simulations indicate that only a moderate amount of cementing agent (< 5 vol%) is required to achieve such a thermal inertia [Piqueux and Christensen, 2009]. Duricrust has a higher thermal inertia than the particles that constitute it, as its thermal inertia is consistent with particle sizes in the range of 0.6-3 mm [Presley and Christensen, 1997]. This is because cementation increases the contact area between grains and, hence, the effective thermal conductivity (and to a lesser extent, the density and heat capacity).
It is difficult to uniquely identify this surface component in
remotely sensed thermal inertia and albedo however, as the thermal inertia values are also consistent with coarse sand (1-3 mm) and the albedo varies with the fraction of bright dust present. The range of 350-600 tiu, comprising coarse sand, duricrust, and a mixture of pebbles and sand (Table 7-3), encompasses a large range in albedo with four peaks at 0.13, 0.22, 0.27 and 0.32 (Figure 7-6). These modes are consistent with clean coarse sand/duricrust (albedo 0.13), and dusty coarse sand/duricrust beneath a thin layer of bright dust of varying spatial coverage (albedos 0.22 and 0.27).
Figure 7-6: Histogram of pixels within the thermal inertia range of 350-600 tiu and their corresponding albedo values. These values are consistent with pure duricrust, coarse sand, pebbles mixed with sand, ice-cemented soil, and mixtures of those components with dust.
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Bridging Chapter: The Influence of Surface Materials on the Subsurface Environment
Table 7-3: Characteristics of end-member martian surface materials. Material
Thermal inertia (tiu) 20-150
Albedo
Emissivity [wavelength]
Dust
Dimensions (grain size)* 2-60 µm
> 0.27
0.92 [7.1-7.4 µm]; ~ 1 at longer wavelengths
Sand
60-2000 µm
150-400
< 0.15
Duricrust
Grains 100300 µm Bulk 0.6-3 mm
350-600
< 0.15
Granules/ very coarse sand
2-4 mm
400-800
< 0.15
Pebbles/ rocks/ bedrock
4-250 mm
800-2068
< 0.15
Boulders/ icecemented soil
> 250 mm
> 2068
< 0.1
> 2500
> 0.4
Ice
References**
[Christensen, 1986a, 1986b; Jakosky, 1986; Kieffer et al., 1973; Presley and Christensen, 1997; Ruff and Christensen, 2002] 0.85-0.95 [9,20 µm] [Edgett and Christensen, 1991, 1994; Mellon et al., 2008b] 0.92-0.97 [20,7 µm]; [Christensen, 1982, greater than dust at 1986b; Edgett and short wavelengths Christensen, 1991; Jakosky and Christensen, 1986; Kieffer et al., 1973; Moore and Jakosky, 1989; Presley and Christensen, 1997] 0.92-0.97 [20,7 µm]; [Christensen, 1982, greater than dust at 1986b; Edgett and short wavelengths Christensen, 1991; Kieffer et al., 1973; Mellon et al., 2008b] 0.92-0.97 [20,7 µm]; [Bell et al., 2004b; greater than dust at Christensen, 1982, short wavelengths 1986b; Golombek and Rapp, 1997; Grant et al., 2004; Nowicki and Christensen, 2007] 0.92-0.97 [20,7 µm]; [Christensen, 1982, greater than dust at 1986b; Golombek short wavelengths et al., 2003; Sizemore and Mellon, 2006] 0.55-0.75 [0.3-3 µm] [Kieffer et al., (CO2 ice) 1976; Mellon et al., 2009a; Paige and Ingersoll, 1985; Putzig et al., 2010]
* Grain size classified according to the Wentworth scale [Wentworth, 1922]. ** The values for each material are derived from a combination of remote sensing and laboratory data.
193
7.1.5.d.
Rock
The term ‘bedrock’ indicates the un-fragmented native rock layers that predominantly underlie, and are the source for the martian regolith.
Where
regolith is absent, for example along steep cliffs and in regions where strong winds dominate and remove small particles, bedrock can sometimes be observed. Thermal inertia values of exposed bedrock are > 1000 tiu, measured in-situ by Opportunity [Christensen et al., 2004b; Cushing et al., 2009]. On Mars, it is more common for bedrock to be overlain by the products of its fragmentation and erosion (such as rocks, sand and dust particles) and to form a small fraction of the TES field of view and the martian surface in general [Edwards et al., 2009]. The term ‘rock’ on Mars has been used to describe fragments of consolidated material of any size > 3 cm in diameter [Golombek and Rapp, 1997; Grant et al., 2004]. Values of thermal inertia for rocks with diameters of 0.03-0.26 m follow the power law: Equation 7-9:
I ~ 5850 × d0.75
where the diameter d is expressed in metres [Golombek et al., 2003]. Martian rocks thus can be expected to have thermal inertias between 420 and 2130 tiu. From Equation 7-9 thermal inertia values > 1000 tiu indicate that the surface is dominated by rocks > 0.1 m diameter or bedrock. Albedo values of martian bare rock are generally < 0.1 [Christensen, 1986b]. Therefore, the combination of high thermal inertia values with low albedo is taken to unambiguously represent exposed bedrock or rocky surfaces. This is illustrated in Figure 7-7, as a peak in albedo at ~ 0.11 is observed for thermal inertia values between 1000-3000 tiu, corresponding to large clean rocks and possibly bedrock. The higher albedo peak corresponds not to rock but to ice, with a potential contribution from bright dust lowering the albedo. Ice-cemented soil may also be included within this range, and may contribute to the higher albedo peak where associated with bright surface frost.
194
Bridging Chapter: The Influence of Surface Materials on the Subsurface Environment
Figure 7-7: Histogram of pixels within the thermal inertia range of 1000-3000 tiu and the corresponding albedo values. These values are consistent with pure rock and bedrock, rock with a mixture of dust, and ice.
7.1.5.e.
Ice
On Mars, water ice and ice-cemented substrate typically have a thermal inertia comparable to boulders and bedrock, with a value of 2506 tiu measured at the Phoenix landing site. This increases the difficultly in discriminating rock and ice in this parameter alone [Mellon et al., 2009a; Putzig et al., 2010] (see Figure 7-7, high albedo peak). The TES albedo of bright clean water ice on Mars is ~ 0.45 as observed in the perennial water polar caps [Calvin and Titus, 2008; Rodriguez et al., 2007]. Dirty water ice, with a moderate < 10 % fraction of dust, has a much lower albedo of ~ 0.23 [Langevin et al., 2005b; Titus et al., 2003]. Furthermore, dirty water ice is also consistent with pure CO2 ice, which has an albedo of ~ 0.32 [Kieffer et al., 2006]. At high latitudes, polewards of ~ 55˚, water ice exists in at least the top 2 m of substrate mixed with regolith and rocky material [Boynton et al., 2002b] (5.2.3), and hence a thermal inertia of the order of ~ 1000 tiu can reflect either bedrock or ice-cemented soil [Bandfield, 2007; Titus et al., 2003]. At low latitudes (< 50˚), water ice is typically insulated by low thermal inertia materials [Smith et al., 2009b], and hence the surface has a much lower thermal inertia, ~ 500 tiu, which is also consistent with coarse sand or a mixture and sand and pebbles. Therefore, water ice can only be unambiguously distinguished if it is clean, exposed, and dominates a large fraction of the pixel area.
195
7.2.
Illustration: Estimating the depth to the melting isotherm
7.2.1. General trends Even without using a sophisticated energy balance model for calculating the surface boundary condition [Mellon et al., 2004b], the temperature range experienced by different materials at the surface (οT) can be estimated by other means.
Once οT is constrained, Equation 7-6 can be applied to model how
temperatures vary at depth. Figure 7-8 is a schematic of the range of subsurface
temperatures experienced by surface materials at 30˚ latitude on Mars across four seasons. The broad range of temperatures experienced on the ground is fixed for each material at 160-240 K, and corresponds to a typical maximum and minimum surface air temperature at 30˚ latitude. This range is a gross overestimate of the actual temperature variations at the surface, as surface materials have very different thermal properties to air. The thermophysical properties used for each material are given in Table 7-4 and the length of a martian year is given in Table 7-5.
Although the orbital thermal inertia measurements cannot reliably be
extrapolated outside of their sensitivity depth of ~ 2 diurnal skin depths, the thermal inertia of the surface materials determines the downwards conduction of solar heat, which dominates the subsurface temperatures within ~ 4 annual skin depths.
The deep subsurface geothermal gradient which determines the
temperatures beneath this depth, is calculated from the Equation 3-2 using a heat flow of ~ 15 W/m2 and the thermal conductivities given in Table 7-4. Although these geotherms do not converge in the image, the steep geothermal gradients corresponding to dust and sand occur through only in a thin surface layer, generally < 10 m [Bourke et al., 2006; Christensen, 1986a] (see 6.1.4.c), and convergence would occur beneath this depth.
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Bridging Chapter: The Influence of Surface Materials on the Subsurface Environment
Table 7-4: Model parameters for Mars’ subsurface temperatures. Material
Thermal Density Heat conductivity (kg/m3) capacity (W/m/K) (J/kg/K)
Thermal Skin Typical diffusivity depth thermal 2 (m /s) for one inertia* martian year (m) -9 1.3×10 0.2 28
References
Dust
0.001
1000
800
Sand
0.1
1750
800
7.1×10-8
0.6
374
Duricrust, pebbles
0.3 (up to 2)
1750
800
2×10-7
2
648
[Murphy et al., 2009; Piqueux and Christensen, 2009]
Icecemented soil
2.5
2018
1040
1.2×10-6
0.8
2294
2.5 (generally 1.5-4.5)
2900
800
1×10-6
3.1
2408
3.4 (-110 ˚C); 2 (0 ˚C)
928
1310
1.3×10-6
5.0
2044
[Mellon et al., 2004a; Sizemore and Mellon, 2006] [Clauser and Huenges, 1995; Mellon and Phillips, 2001; Sizemore and Mellon, 2006; Turcotte et al., 2002] [Clauser and Huenges, 1995; Mellon et al., 2009a; Titus et al., 2003]
Rock
H2O ice
*Calculated from Equation 7-7
[Mellon and Phillips, 2001] [Heldmann et al., 2005; Mellon and Phillips, 2001; Murphy et al., 2009]
197
Despite its simplicity, this basic schematic is useful for illustrating the shape of temperature fluctuations at depth. Figure 7-8 reveals the relative temperature environments of fine grained materials which rapidly dampen annual variation within 1-2 m depth (dust and sand), and coarse grained and cemented materials which conduct temperature variation to 5-10 m depth (pebbles, duricrust, rock and ice). To compare the subsurface temperature environment with the stability of water and the likelihood of brines, a more sophisticated approach is required. Specifically, more accurate surface boundary conditions must be implemented for each material, to reflect their actual temperature variations at the surface. This range will be albedo and thermal inertia dependent. As discussed in 7.1.3, high albedo materials reflect more of the incident solar radiation, absorb less heat, and hence experience a smaller range of temperatures. A difference of 0.2 in albedo (between 0.15 and 0.35) can result in a difference of ~ 20 K in the maximum temperature experienced by the material [Mellon et al., 2000]. Similarly, the higher the thermal inertia of the material, the smaller the amplitude of its diurnal temperature variations, as the material is effectively dispersing the incident heat throughout its thermal bulk. Below an alternate approach to constraining the surface boundary condition is presented, providing a more accurate measure of surface and therefore subsurface temperatures. Table 7-5: Length of martian time periods (in seconds) Day 8.9 x 104
Northern Seasons Summer Autumn 1.6×107 1.3×107 (179 sols) (143 sols)
Year Winter Spring 1.4×107 1.7×107 (154 sols) (193 sols)
5.9× 107
*Average season is 1.5 x 107s; ߜ௬ ൌ ʹ ൈ ߜ௦௦ ൌ ʹͷǤ ൈ ߜௗ௬ where ߜis skin depth from Equation 7-3, and is proportional to the square root of the period.
7.2.2. Example of Gale Crater Gale crater provides an appropriate site to illustrate the role of surface materials in determining the subsurface thermal environment, as it is the landing site of the Mars Science Laboratory in August 2012 [Golombek et al., 2011] (Figure 7-9) and its surrounding surface materials have been significantly studied and characterised. Gale crater is a 150 km diameter impact crater. It has strong astrobiological significance demonstrated through the layered deposits in its central mound which show concentrations of sulphates and phyllosilicate clays,
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which were most likely precipitated out of water [Le Deit et al., 2011; Milliken et al., 2010]. The central mound also displays alluvial fans, inverted channels and other morphologies which support the interpretation of it having strong fluvial modification [Thomson et al., 2011]. Multiple fluvial channels also flow into the crater and thus liquid water has likely interacted with the floor materials that the Curiosity rover will be sampling [Anderson and Bell III, 2010; Pelkey et al., 2004].
Figure 7-8: Estimated maximum range in subsurface temperatures during a martian year at ~ 30˚ absolute latitude. The freezing temperatures of average Mars salinity water (at -10 ˚C) and maximum liquid range (at -100 ˚C) are shown. The skin depth encompassing the seasonal solar variation within a martian year is shown in brackets for each material, and is the depth at which the subsurface temperature fluctuations drop to 1/e of their surface value. The skin depths are given in Table 7-4 and were calculated in Table 7-5. Below several seasonal skin depths, the geothermal gradient dominates subsurface temperatures and is determined by the ratio of the heat flow to the thermal conductivity of the material.
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Figure 7-9: Mosaic of Gale Crater in daytime visual imagery from THEMIS, centred at 5.3 ˚S, 137.7 ˚E. Crater diameter is 150 km. Image credit: NASA/JPL-Caltech17.
The Thermal Emission Imaging System (THEMIS) onboard Mars Odyssey measures surface reflectance in multiple red and infrared bands [Christensen et al., 2004c], with nighttime (AM) infrared measurements centred at 12.58µm (band 9) enabling surface temperature values to be derived [Christensen et al., 2003]. Within the metadata of a band 9 THEMIS-IR image is the maximum and minimum surface temperature and best-fit thermal inertia across all pixels at the time the image was taken (image metadata in Appendix A, Table A3). 211 band 9 THEMIS-IR images covering the Gale Crater region across all four martian seasons are considered. The maximum and minimum surface temperatures within the images can be used to derive the surface boundary condition for each end-member material. Given that all surfaces within the region will experience a similar insolation environment (with some variation due to slopes), then differences in surface temperatures can
17
http://apod.nasa.gov/apod/image/1107/GaleCraterMars_themis.jpg
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Bridging Chapter: The Influence of Surface Materials on the Subsurface Environment
be attributed to thermal inertia and albedo differences. To classify the thermal inertias of the images as surface materials, Table 7-3 is used. The classification thresholds are summarised in Table 7-6. Figure 7-10 shows the relationship between the minimum and maximum thermal inertia and surface temperature across each of the 211 images. All images are taken at a local time between 3-5 am (Table A3). In the early martian morning, low thermal inertia materials will have lost the solar heat which warmed them during the previous day, either through horizontal heat conduction to adjacent materials or radiation to the cool martian atmosphere. Hence they will be the coolest materials on the surface, and will be responsible for the lowest surface temperature in the image [Ditteon, 1982].
In contrast, high thermal inertia
materials have a higher capacity to store heat and will be the warmest materials on the surface during 3-4 am [Michalski and Fergason, 2009]. As expected these trends are observed in Figure 7-10.
201
Figure 7-10: Correlation between thermal inertia and surface temperature at Gale crater from THEMIS nighttime infrared images. All images were taken between in the early morning between 3-5 am (Appendix A, Table A3). Clearly this approach constrains some materials’ temperature variability better than others. For example, duricrust will only be the lowest temperature material within a THEMIS image when sand and dust are absent. If sand and dust dominate at least one pixel within an image, then the minimum image temperatue will be due to these materials, and no information will be available on the minimum temperature of duricrust.
The ability to constrain the full range of annual
temperature variability in each material therefore depends on their distribution on the surface. Another caveat to this approach is the use of 3-5 am temperatures
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across four martian seasons as a proxy for the annual temperature variation of the surface. This limitation results in a lower estimate of the maximum temperature of the surface, particuarly for low thermal inertia materials. At the latitude of Gale crater, the maximum temperatures during summer are not reached until noon [Kieffer, 1976].
In reality, each surface material will experience a larger
temperature range than is derived from Figure 7-10, with the real upper boundary being several degrees warmer than that given in Table 7-6. Despite its limitations, this technique of deriving the range temperatures experienced by each material at the surface from the metadata in THEMIS infrared images, and applying the heat conduction model of Equation 7-6, will be shown below to provide realistic trends in the temperature variability of surface materials and useful information on the subsurface temperature environment. Table 7-6: Properties of Gale crater materials from THEMIS-IR band 9 imagery taken across four seasons. Interpretation of dominant surface material Dust Sand Coarse sand, pebbles, duricrust Rock Ice-cemented soil Water frost *
THEMIS thermal inertia range (tiu) 20-100 100-300 300-800
THEMIS morning temperature range across four seasons (K) 148-180 148-190 175-220
800-2000 800-2000 800-2000
190-220 190-220 190-220
Density Heat Annual (kg/m3) capacity skin (J/kg/K) depth (m)* 1000 800 0.11-0.54 1750 800 0.31-0.93 1750 800 0.93-2.5 2900 2900 928
800 800 1310
1.5-3.7 1.5-3.7 2.85-7.13
P = 5.9 ×107s (Table 7-5)
To model subsurface temperatures, the temperature ranges derived from the THEMIS images (summarised in Table 7-6) were used to provide the surface boundary condition for each end-member surface material (identified from the THEMIS thermal inertias).
This provided an estimate of οT for dust, sand,
duricrust, pebbles, rocks, and ice.
In combination with the modelled
thermophysical properties of each material (skin depths in Table 7-6, calculated from Equation 7-3), the heat conduction model of Equation 7-6 was then be applied. The results are shown in Figure 7-12, providing an estimate of the varied subsurface temperature environments present in the Gale crater region within a martian year. In general, high thermal inertia materials such as duricrust, rock and
203
ice experience less surface temperature variation within a year but have a higher maximum and minimum. Due to their effective heat conduction, they provide the warmest subsurface environment below 0.5 m. Low thermal inertia materials experience a much a wider fluctuation in temperature however dust, which has the lowest thermal inertia, does not experience the largest οT at the surface. This is due to the effect of albedo. Dust has a high albedo (Table 7-3) which reduces its absorption of solar heat and decreases its temperature variability [Smith, 2004]. In general, dust and sand effectively insulate the subsurface below 50 cm depth from the seasonal variations at the surface.
Figure 7-11: Temperature variability across four seasons within materials in the Gale crater region. From Figure 7-11 the subsurface temperatures within all materials at Gale Crater allow for some period of water ice stability, as all materials experience temperatures below 200 K (the frost point; 5.1.1). Duricrust, rocks and ice would likely accumulate water frost transiently above ~ 5 m depth during high humidity and low temperature, and duricrust could be permanently ice cemented between 5-50 m depth. Dust and sand could have permanent concentrations of ice, with the
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Bridging Chapter: The Influence of Surface Materials on the Subsurface Environment
volume varying with humidity. These latter materials occur frequently in Gale Crater and hence will be observed by MSL (Figure 7-12).
Figure 7-12: Variation in grain size at Gale Crater from THEMIS thermal inertia measurements [Hobbs et al., 2010] (Fig. 4). Liquid water is expected to be minimal in this region. Thin films would likely coat all materials during high humidity, with the highest thin film content occurring in the warmest materials during the evening humidity maximum [Carrozzo et al., 2009; Searls et al., 2010]. Hence, the highest volume thin films would be found coating rocks and duricrust and on the surface of water ice, in the evening to early morning. Perchlorate brines analogous to that observed at the Phoenix site [Hecht
205
et al., 2009] (5.3.3) could potentially occur during summer within the top ~ 3 m depth in association with water ice, rock and duricrust, if the salts are present. There is growing evidence that these salts may be widespread [Clark et al., 1976; Navarro-González et al., 2010; Toulmin et al., 1976]. Larger volumes of liquid brines with eutectic > -10˚ (average martian brine; 6.1.1) would not occur in the shallow regolith. In summary, although water ice is expected in dust and sand (at least during the nightly high humidity), the most likely location for subsurface liquid water in the region of Gale crater is transiently during the evening in duricrust and rocks. This liquid would exist either as thin films or deliquescent brines sourced from atmospheric water vapour. MSL however will be unlikely to sample these latter materials as they occur around the rim of Gale Crater (Figure 7-12).
7.3.
Summary
The examples presented in this chapter highlight the significance of thermal inertia and albedo in describing the subsurface temperature environment and, most importantly, the likelihood and stability of shallow water. In Chapter 8, maps of thermal inertia and albedo are utilised to map the distribution of materials across the martian surface and assist in constraining the global subsurface water environment.
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8. REMOTE SENSING OF THERMAL MATERIALS ON THE MARTIAN SURFACE
‘Someone once told me that time was a predator that stalked us all our lives. But I rather believe that time is a companion who goes with us on the journeyreminds us to cherish every moment, because they'll never come again. What we leave behind is not as important as how we've lived. After all, Number One, we're only mortal.’ - Captain Jean-Luc Picard, ‘Star Trek Generations’
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In Chapter 6, the presence of shallow environments on Mars within the temperature and pressure regime of liquid water was identified. At low latitudes with mean annual temperatures above ~ 250 K, these shallow water environments could potentially be hospitable to life during their diurnal liquid availability. Shallow subsurface habitability is strongly constrained by the thermal properties of the surface materials, that determine how insolation is distributed to the subsurface and hence the shallow temperature variability. Therefore, the surface materials establish the occurrence of near-surface P-T conditions within the metastability regime of liquid water and the temperature thresholds of Earth life. The general trends shown in Chapter 7 were that high thermal inertia and low albedo materials, such as rock, effectively transfer solar heat to the shallow regolith. However if temperatures are already high (such as near the equator), fine-grained dust with low thermal inertia and high albedo most effectively insulates the subsurface, allowing it to experience more stable temperature conditions centering around the warm seasonal mean. In this chapter, the first requirement of a heat-conduction model for assessing the habitability of the shallow martian subsurface is addressed, which is to estimate thermal behaviour of materials on Mars and their occurrence on the surface. Thermal inertia and albedo provide key information on the distribution of surface materials on Mars. These parameters have been mapped globally on Mars by the Thermal Emission Spectrometer (TES) onboard Mars Global Surveyor, enabling the spatial occurrence of two-dimensional groupings of pixels in global thermal inertia and albedo to be studied.
These two-dimensional groups are termed
‘thermophysical units’, in reference to the influence of thermal inertia and albedo on the temperature variations within a given material.
The units reflect the
dominant materials on the surface. A number of authors have defined and mapped the distribution of thermophysical units [Mellon et al., 2000; Paige and Keegan, 1994; Paige et al., 1994; Palluconi and Kieffer, 1981; Putzig and Mellon, 2007a; Putzig et al., 2005].
In this chapter, a more automated, less deterministic,
algorithmic classification methodology for defining thermophysical units is employed, providing a new and more detailed mapping of martian surface materials and their thermal behaviour.
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8.1.
Background
8.1.1. Previous maps of martian thermophysical units Mapping of remotely measured thermal inertia and albedo provides insight into the nature of surface materials and near-surface geology. Early low resolution mapping of thermal inertia and albedo from the Viking Infrared Thematic Mapper (IRTM) enabled the determination of two dominant thermophysical units on the martian surface. These units were characterised by: low thermal inertia and high albedo, reflecting martian dust (unit A); and high thermal inertia and low albedo, reflecting grains from sand-size to pebble-size with a high rock fraction (unit B) [Palluconi and Kieffer, 1981].
Higher resolution thermal inertia and albedo
measurements provided by TES allowed a third unit to be defined, due to the addition of an intermediate albedo peak in the global histogram.
This unit
reflected a similar range of grain sizes to unit B but was interpreted as indurated fines (duricrust) [Mellon et al., 2000]. The latter interpretation is consistent with visual and radar observations of the surface [Christensen and Moore, 1992], although it is also plausible that unit C was a unit B surface with a ubiquitous covering of dust.
As the spatial coverage of TES data became more extensive and the ‘lookup model’ for deriving thermal inertia more sophisticated (allowing the derivation of higher values of thermal inertia more representative of the surface), several new units were mapped. The next mapping derived 7 thermophysical units shown in Figure 8-1 [Putzig et al., 2005]. Units A-C closely matched previous maps, however the additional units (D-G) provided greater detail on intermediate thermal inertia and albedo (D and E) as well as extremely high thermal inertia (F) and extremely high albedo (G). Two significant differences separated this latest work from previous thermophysical mappings [Putzig et al., 2005]. The first, was the incorporation of units distinguished by a thermally (less than the ~ centimetres sensing depth of thermal inertia) or optically (less than the ~ microns sensing depth of albedo) thin surface covering. For example, unit G in Table 8-1 is interpreted as a unit A surface covered by dust only microns thick, and therefore dominating the albedo but not the thermal inertia of the pixels. The second significant difference between Putzig
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Remote Sensing of Thermal Materials on the Martian Surface
et al.’s [2005] map and previous works was the incorporation of a new unit with low thermal inertia and low-intermediate albedo dominating the southern latitudes > 65˚S. With thermal inertias consistent with predominantly dust sized grains, the low albedo values included within the class led to an interpretation of low density surface materials from desiccation of an ice rich surface layer.
Figure 8-1: Thermophysical map of Putzig et al. [2005] (Fig. 5) showing the division of the martian surface into 7 thermophysical units, described in Table 8-1. Although this explanation is consistent with the high abundance of ice morphologies viewed on unit D terrain, it presents several problems: -
Ice related morphologies such as polygons also occur extensively in Utopia Planitia [Clifford and Parker, 2001; Pearce et al., 2011; Soare et al., 2011] within units B and C. The occurrence of these morphologies in class D therefore provides an insufficient explanation for the thermal inertia values within the class.
-
Typical hydrogen content between 70-80˚ in both the northern and southern polar terrain is similar, indicating that concentration of H2O ice within the soil varies between 20-35 wt% in the top 2 m [Maurice et al., 2011] (Figure 5-4). This is inconsistent with a significant substrate and
211
density difference between the shallow subsurface at each pole that is indicated by the map [Putzig et al., 2005]. -
From the global histogram of thermal inertia and albedo (Figure 4 of [Putzig et al., 2005]), the low albedo values < 0.2 incorporated into class D correspond to thermal inertia values > 100 tiu and hence could potentially be explained by fine dark sand consistent with observations by landers on the surface [Kieffer et al., 1973; Presley and Christensen, 1997].
Once these lower albedo, higher thermal inertia pixels are
separated out from class D, the remaining surfaces can be more simply described as a mixture of bright dust with a small fraction of sand at the sub-pixel scale. This latter point on more than one component material within a thermophysical class can be addressed by the alternative methodology applied in this work, described in 8.1.2 below. Table 8-1: Extreme values of thermophysical units from Putzig et al. [2005] (Tab. 1).
All maps of martian thermophysical units show a trend of fine dust-sized grains being dominant in the low latitudes, coarse sand in the mid-latitudes, and ice at the high latitudes [Mellon et al., 2000]. This is consistent with visual observations of the surface [Christensen, 1986b].
The mechanical and chemical processes
responsible for the global gradient in dust coverage and grain size are not examined here but are a combination of both aeolian transport of fine particles [Greeley et al., 2006, 2008], weathering by ancient
volcanic and fluid flows
[Gooding et al., 1992; Squyres et al., 2004a, 2004b], and other weathering processes such as the abrasion of large grains by smaller grains, and chemical interactions with liquid water and water ice [Grindrod et al., 2012; McGlynn et al., 2012]. The north and south polar terrain have thermal inertias that differ by ~ 200 tiu. This is
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due in part to errors in processing thermal inertia values at the poles [Putzig and Mellon, 2007a; Putzig et al., 2005] (see 8.2.1.a) but also due to the extensive coverage of transverse dunes around the north polar residual cap that have been observed to have a high fraction of ice [Paige et al., 1994; Putzig et al., 2010]. In contrast, the surface of the south polar layered deposits consists of finer dark grains [Paige and Keegan, 1994; Vasavada et al., 2000]. Many impact structures show higher thermal inertias around their rims and on their impact ejecta, compared to the surrounding terrain [Betts and Murray, 1993; Mellon et al., 2000]. Hence, some impact structures larger than the 3×3 km area of TES pixels should be recognisable through the occurrence concentric circular structures in a distinct thermal inertia, delineating the crater rim and its ejecta blanket.
Although
previous thermophysical maps may be consistent with this, there have been no published examples.
All thermal inertia values occur over a wide range of
elevations and thus a correlation between thermophysical units and elevation is not observed [Palluconi and Kieffer, 1981; Putzig et al., 2005].
8.1.2. Motivation for a different methodology All previous thermophysical mappings applied a similar methodology to determine divisions between thermal inertia and albedo units [Mellon et al., 2000; Palluconi and Kieffer, 1981; Putzig et al., 2005]. These works employed a deterministic method of manually applying thresholds around the globally dominant values of thermal inertia and albedo. For example, the primary classes A, B and C of Table 8-1 encompass the three peaks in albedo in the global histogram (Figure 7-2). This technique is sensitive to the global modal values and hence it enables the surface components of bright fine dust, dark fine sand and bright ice to be clearly distinguished, due to their extreme values of albedo and thermal inertia and extensive spatial coverage (Table 7-3). Each pixel in the TES thermal inertia and albedo maps encompasses a surface area of order 9 km2 and hence the majority of pixels will be not be homogenous at this scale. Instead, the majority of pixels will be ‘mixels’. Mixels refers to pixels that are not occupied by a single category of surface materials as instead they have significant sub-pixel variation [Fisher, 1997; Muad and Foody, 2012; Puyou-Lascassies et al., 1994; Zhang and Foody, 1998]. This is due to the resolution of the pixel not being fine enough to capture a sufficient level of detail on the ground. Such pixels will have orbital values that correspond
213
to a linearly weighted average of the components within the pixel, with the weights reflecting the fraction of the pixel encompassed by each component. Therefore, although the methodology applied by previous authors in defining the unit boundaries enables a clear detection of large homogenous martian surfaces, the majority of pixels display a range of grain sizes with varying albedo and are difficult to manually separate from the global variation. This is illustrated in Table 8-2. Table 8-2: Unambiguous interpretation of surface materials. TI threshold
Albedo threshold
1000-3000 500-800
< 0.15 < 0.15
150-400 1000
< 0.15 > 0.25 > 0.3
Surface material classification
Fraction of map pixels % Rock 0.06 Duricrust/pebbles/coarse 0.10 sand Sand 19.10 Dust 12.77 Ice 1.29 Total: 33.32
Based on the response of martian surface materials in thermal and visual wavelengths (discussed in 7.1.5), there are some values of thermal inertia and albedo that allow a pixel to be uniquely interpreted as one of the end-member surface materials (with the exception of duricrust). Table 8-2 reveals that a unique mapping from the thermophysical response of the surface to the dominant surface materials only occurs at extreme values of thermal inertia and albedo. Only ~ 30 % of the global map can be unambiguously interpreted as a > 2 cm surface layer of dust or > 5 cm layer of sand (thickness based on skin depth; Equation 7-3), with ~ 5 % corresponding to larger grain sizes (rocks, ice). The remaining approximately 65 % of the map, corresponding to the two lower modes in global albedo and the higher mode in thermal inertia in Figure 7-2, is mixels. Given these observations, there is scope to apply the less deterministic methodology of algorithmic classification to determine the boundaries between thermal inertia and albedo units. The aim of this method is to provide greater detail and a higher level of accuracy in the spatial distribution of martian surface materials. In particular, this work aims to provide a more sophisticated treatment
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Remote Sensing of Thermal Materials on the Martian Surface
of mixels so that the relative fractions of thermally distinct surface materials within heterogeneous surfaces can be constrained.
8.2.
Methodology
The framework for interpreting thermal inertia and albedo values as materials on the martian surface, combining direct sampling from landers, experimental data on martian analogue materials, and theoretical data on the physical behaviour of grains under martian conditions, was presented in Chapter 7. The datasets utilised in this work and the algorithmic classification techniques are detailed below.
8.2.1. Observations 8.2.1.a.
Thermal inertia & albedo
The technical specifications of Mars Global Surveyor’s Thermal Emission Spectrometer (TES) used to collect the visual and infrared data from which albedo and thermal inertia are derived are described in detail in Christensen et al. [2001a]. Information on the datasets was given in 7.1.2 and 7.1.3. TES covers multiple wavelength ranges: from 0.3-2.7 µm (a single visible band used for albedo), 5.5100 µm (a single infrared band used for thermal inertia), and a multiband thermal infrared spectrometer. The spatial resolution of each sensor is ~ 3 km. Mapping in the visible and infrared band was undertaken during 1999-2001 (Mars Years 2426 [Clancy et al., 2000]).
Both albedo and thermal inertia datasets have
dimensions of 7200 × 3600 pixels corresponding to the intrinsic sensor resolution of 0.05˚ per pixel (~ 3×3 km pixels). Real data constitutes ~ 93 % of the thermal inertia map and ~ 35 % of the albedo map [Putzig and Mellon, 2007a](Putzig pers. comm.). Gaps in geographic coverage occur in both the albedo and thermal inertia maps are due to the orbital tracks, spacecraft calibration, limb observations that were rejected, and surface brightness temperatures outside the range of the thermal inertia lookup model [Putzig, 2006] (Equation 7-8). The remainder of the data in the map has been bilinearly interpolated between the observations. Although the albedo map has a low fraction of real data, it has a large overlap with the time period during which the infrared mapping used to derive thermal inertia occurred (discussed below).
The thermal inertia dataset used here is more
complete than that used by previous works [Putzig et al., 2005], where real data
215
constituted only 60 % of the map [Putzig, 2006]. Furthermore, from the discussion in 7.1.4, the orbital values of both albedo and thermal inertia match well with the values of the fine-component of the surface measured by landers. The albedo measurements were taken within a single martian year (MY24). This year was selected as it had minimal localised dust storm events [Cantor et al., 2002] and generally a lower dust optical depth (the atmosphere was more transparent) than MY25 and MY26 [Smith, 2004; Tamppari et al., 2008]. Thus the albedo values in MY24 will be more representative of the mean surface materials and less affected by scattering due to atmospheric dust.
The albedo change
between MY24 and MY26 was less than ±0.06 over the vast majority of the martian surface [Putzig and Mellon, 2007a]. The instrument uncertainty in albedo values is approximately ± 0.01 [Christensen et al., 2001a]. The albedo map has not been corrected for roughness which can alter the effective albedo of the surface determined from orbit by forward or backward scattering, with the effect increasing with incidence angle [Pleskot and Miner, 1981; Squyres and Veverka, 1982]. From mapping of the surface roughness at ~ 1 m vertical resolution, this photometric scattering affect will be highest for surfaces around Olympus Mons, Tempe Terra, Valles Marineris and Olympia Planum [Neumann et al., 2003]. Data used to produce the 2007 thermal inertia dataset is taken predominantly over MY24-25 with a small fraction of the extended mapping included from MY26 [Putzig and Mellon, 2007a]. Nighttime thermal inertia values were used due to the uncertainty associated with each derived value estimated to be < 10 % [Putzig et al., 2005] and hence lower than the 17 % upper bound for daytime values. These uncertainties are a combination of instrument measurement error, and uncertainties in the thermal model and the interpolation scheme used to derive thermal inertia [Putzig, 2006]. Nighttime temperature measurements were taken at 2 am. Atmospheric corrections were made to the brightness temperatures, by subtracting the maps of representative atmospheric dust conditions and water ice clouds opacities from the orbital reflectance [Mellon et al., 2000]. Pixels with particularly high opacity from dust or clouds during the mapping phase were rejected [Smith, 2004]. Approximate slope corrections were made [Putzig and Mellon, 2007a] to account for the changes in insolation experienced on steep
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Remote Sensing of Thermal Materials on the Martian Surface
slopes and the variation in atmospheric-path-length [Putzig and Mellon, 2007b]. The lookup model used to derive thermal inertia excluded surfaces at and below the frost point of CO2 [Putzig et al., 2005] (≤ 160 K). The final thermal inertia values are the medians of 36 seasonal maps of data obtained across the 3 martian years, extending from ± 87˚ (due to the orbital inclination of the spacecraft). This thermal inertia dataset differs from that utilized by previous works [Putzig, 2006; Putzig et al., 2005], as the earlier model for deriving thermal inertia only computed values within the range of 0-800 tiu [Putzig, 2006]. Thermal inertia values > 800 tiu encompass only 5.7 % of the newer 2007 map and so only a small fraction of pixels have values outside the earlier (2005-2006) lookup model. Additionally, the higher geographic coverage of the 2007 thermal inertia map introduces values that may differ from the interpolated values in the earlier maps. The thermophysical map derived in this work is compared to the most recent previous mapping of Putzig et al. [2005] in 8.3.3. The sensing depth of TES in infrared wavelengths (relevant to thermal inertia) is on the order of centimetres [Christensen et al., 2001a] (~ 2 cm for dust, 4 cm for coarse sand, Table 7-2) and is on the order of microns for visual wavelengths (relevant to albedo). The actual sensing depth of thermal inertia is determined by the diurnal thermal skin depth of the surface materials (Equation 7-3). The range of diurnal skin depths for martian surface materials is ~ 2-30 cm [Christensen and Moore, 1992] and hence TES provides information only on the surficial levels of the crust.
8.2.1.b.
Other data
Visual imagery from the High Resolution Imaging Science Experiment (HiRISE) and the Thermal Emission Imaging System (THEMIS) were utilised to validate the interpretation of surface materials from the thermophysical map. The HiRISE images provided the surface reflectance in the red wavelength between 570-830 nm at 25 cm/pixel [McEwen et al., 2007b].
The THEMIS images provided the
surface reflectance/emission in the infrared wavelength band centred at 12.57 µm (width ~ 1 µm) at 100 m/pixel [Christensen et al., 2004c]. The details of these images are given in Appendix A. Map-projected HiRISE images were accessed from http://hirise.lpl.arizona.edu//.
High-quality band 9 THEMIS images were
217
identified using the Java Mission-planning and Analysis for Remote Sensing system (JMARS: http://jmars.asu.edu/). Table 8-3: Global datasets for Mars. 1
Dataset Albedo (MY24)
2
Nighttime thermal inertia (MY24-26)
3
Dust cover index Elevation
4
5
6
Putzig’s thermophysical units Dunes
7
Geology
Source http://lasp.colorado.edu /inertia/2007/albedo.ht ml http://lasp.colorado.edu/ inertia/2007 ftp://pdsimage2.wr.usgs. gov/pub/pigpen/mars/te s/putzig_thermal_inertia/ http://www.mars.asu.ed u/~ruff/DCI/dci.html ftp://pdsimage2.wr.usgs. gov/pub/pigpen/mars/m ola/ Nathaniel Putzig (pers. comm. February 2010)
Ref. [1]
Resolution 1/20˚, ~ 3 km
Extent 80˚S – 80˚N
[1]
1/20˚, ~ 3 km
80˚S – 80˚N
[2]
1/16˚, ~ 3.7 km 1/128˚, ~ 460 m
80˚S – 80˚N
[4]
1/20˚, ~ 3 km
80˚S – 80˚N
http://pubs.usgs.gov/of/ 2010/1170/
[5]
65˚S – 65˚N
http://webgis.wr.usgs.go v/pigwad/down/mars_ge ology.htm
[6]
Polygon areas of dune fields > 1 km2 Polygon areas of geologic contacts
[3]
87˚S – 87˚N
90˚S – 90˚N
[1] [Putzig and Mellon, 2007a]; [2] [Ruff and Christensen, 2002]; [3] [Smith et al., 2001]; [4] Updated version [Putzig, 2006], original publication [Putzig et al., 2005]; [5] [Hayward et al., 2008]; [6] [Skinner et al., 2006]
All other datasets used in this study are shown in Table 8-3. Boundaries of martian dunes have been mapped from THEMIS visible and infrared imagery (100 m/pixel resolution) and Mars Orbiter Camera narrow angle imagery (1.5-12 m/pixel) [Hayward et al., 2007, 2008]. The dune catalogue is likely incomplete for sub-pixel dune fields smaller than 10 km2. 547 distinct dune fields are included. The spectral signatures of martian dust from emissivity between 7.1-7.4 µm has been mapped by TES to provide a dust cover index [Ruff and Christensen, 2002] (7.1.5.a), with a low DCI indicating a high dust coverage. Boundaries of major geologic units and surface morphologies have been mapped globally from Viking Orbiter imagery at 130-300 m/pixel [Skinner et al., 2006]. 92 distinct geologic units are included. The global elevation map is produced from the Mars Orbiter Laser Altimeter which
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Remote Sensing of Thermal Materials on the Martian Surface
mapped topography at ~ 1 m vertical resolution and ~ 460 m/pixel spatial resolution [Smith, 1998; Smith et al., 2001].
8.2.1.c.
Data processing
Data was analysed using commercial GIS and remote sensing software: i.
ArcMap 10 for description, mapping and visualization of the datasets
ii.
ERDAS Imagine for algorithmic classification
iii.
IDL for data manipulation prior to importing into ArcMap and subsequent analysis
iv.
ISIS3 for image processing
v.
JMARS for identifying HiRISE and THEMIS imagery
Datasets 1 and 5 in Table 8-3 were read into an IDL procedure to convert their binary format to ASCII.
They were then read into ArcGIS via the
‘asciitoraster_conversion’ routine. Datasets 2-4 in Table 8-3 were available for direct import as raster layers (grids) into ArcGIS and datasets 6-7 as vector layers (polygons). All maps used equidistant cylindrical coordinates projected to the MARS 2000 IAU projection system [Hare et al., 2006]. Images were processed in ISIS3 using routines freely available from http://isis.astrogeology.usgs.gov/Isis2/ (procedure described in Appendix B).
Pre-processing of the images by the
instrument teams, in the form of radiance calibration, stitching and geometric projection, was undertaken before the images were made available.
8.2.2. Algorithmic classification This research focuses on the application of a different methodology for determining the thresholds of thermal inertia and albedo classes, with the aim of developing a more sophisticated and detailed classification of surface materials. This can in turn provide constraints on subsurface habitability. The algorithmic classification techniques used here have made a valuable contribution to the remote sensing of Earth land cover, and have been shown to provide useful information when operated on two independent datasets (such as 2 bands of an image). They are described in detail below.
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8.2.2.a.
Success in Earth remote sensing
The term ‘algorithmic classification’ refers to a broad suite of algorithms that define multi-dimensional clusters by employing both: one, some measure of similarity to define classes; and two, some formula for class membership to assign pixels. The two algorithms used in this work are ISODATA (to determine initial class boundaries) and the Maximum Likelihood Classifier (to assign pixels to classes). Both were widely used in conjunction in early environmental remote sensing with LANDSAT. One of the earliest presentations of the technique of unsupervised classification was given in the 80’s [Walker et al., 1986]. This work used digital algorithms to define multi-dimensional clusters across four wavelengths of surface reflectance, and interpreted these clusters as distinct land cover features. Algorithmic classification is now widely used in the interpretation of Earth remote sensing data [Hall et al., 1995] and is particularly applied to the classification of surface materials [Belward et al., 1990; Duda and Canty, 2002]. The methodology is applicable to solving a wide variety of problems, in fields as diverse as: discriminating false from real detections of single particles in numerous x-ray images [Yoon et al., 2011]; detecting the most likely geographic origin of different strains of avian influenza [Breland et al., 2008]; characterising galaxy spectra from the Sloan Digital Sky Survey [Almeida et al., 2010]; and linguistic classification of movie reviews as either positive or negative [Turney, 2002]. Algorithmic classification involves using a clustering algorithm to group pixels in parameter space based on a measure of similarity, typically the minimisation of pixel separation within a group. This is done without a priori knowledge of the natural groupings present within the data.
By categorising observations and
minimizing the differences between members in a group, the observations that are most similar can be identified [Duda and Hart, 1973; Murray and Estivill-Castro, 1998; Tou and Gonzalez, 1974]. In the case of physical properties of surface materials (such as grain size), a class can be interpreted as either a pure material, or a mixture of component materials on the surface (mixels). The values for each class are then used to determine some of the attributes of those materials [Jupp et al., 1986; Murthy et al., 2003]. This unsupervised clustering is often necessary with remotely sensed data of a large or complex area as the actual or optimal number of natural clusters in the data is not known. Also when the number of data points is
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large, clusters are a useful way to reduce the data variation by replacing each class with the class mean. Field data can then be collected to define the link between remotely sensed pixel values and the physical interpretation of classes. Such clustering techniques are thus very useful for analysing global datasets of the martian surface but few previous studies have used them (exceptions include the application of classification algorithms to martian topographic data [Bue and Stepinski, 2006; Miller et al., 2011; Stepinski et al., 2006]). Typically the classes provided by an unsupervised classification algorithm (such as ISODATA) must be verified to determine that they are a reflection of real physical properties. This can be done through groundtruthing or examination of visual imagery [Lee and Shan, 2003]. Ideally, a combination of algorithmic classification and real training sites would be used to define the transformation from remote observation (pixel values) to physical properties such as grain size [Dubath et al., 2011]. Frequently large numbers of clusters will be generated surpassing the level of detail that is available from groundtruthing and hence during the classification process they will be reduced (by merging clusters with similar statistics) to a manageable number for interpretation [Aniya et al., 1996]. Some factors, such as a large target area or difficulty in accessing the target area [Gratton et al., 1990], can negate the possibility of sufficient groundtruthing.
In this case, instead of
including real training data in the classification process, a supervised classification algorithm is employed to use the training classes from an unsupervised classifier [Paul, 2001]. Supervised classifiers utilise pre-determined attributes of spectral classes, and a rule for class membership, to assign pixels to clusters [Jia and Richards, 1994, 1999].
For example, the maximum likelihood classifier requires a mean and
standard deviation for each class, and describes the probability of class membership by a multidimensional Gaussian [Richards, 1986]. In the absence of real data on the signatures of the physical classes described by a map (such as the range of surface reflectances corresponding to the land cover type of water), or on the a-priori probabilities of cluster occurrence (such as what fraction of the study area corresponds to water), a supervised classifier can be paired with an unsupervised clustering algorithm [Strahler, 1980; Yang et al., 2010]. In this case,
221
ISODATA or another clustering algorithm passes through the dataset and clusters it probabilistically, simulating a training dataset. The supervised classifier then ‘learns’ from the statistical output of the first to describe the cluster attributes and partitions the dataset into classes using a given decision rule for class membership [Johnson and Wichern, 2007]. The accuracy of algorithmic clustering and classification can be determined using groundtruthing data through the construction of an error matrix. A key feature of an error matrix is to identify the overall accuracy - the percentage of correctly classified pixels. This is examined by comparing the percentage of pixels within a class to the fraction that have been identified as truly having similar land-cover on the ground [Congalton, 1991].
Studies that assessed the overall accuracy of
unsupervised classification paired with the maximum likelihood classifier operated on LANDSAT-TM, MSS and ETM imagery, reported accuracies of 60-90 % [Miller and Yool, 2002; Murthy et al., 2003; Pal and Mather, 2003; Sader et al., 1995].
Hence this technique of combining training data, derived from
unsupervised classification, with a supervised classifier, such as maximum likelihood, can provide spectral clusters that translate into useful information on surface properties. The algorithms used in this study are described below, and the methodology is illustrated in Figure 8-3.
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Remote Sensing of Thermal Materials on the Martian Surface
Figure 8-2: Illustration of ISODATA clustering in a two-dimensional dataset [Richards, 1986] (Fig. 9.2). Two classes are to be found by the algorithm. The points indicate the location of the pixels in multispectral space and m 1,2 mark the class means recalculated with each iteration. The SSE value is the ‘sum of squared error’ which is the sum of the distances between each pixel and the mean of the class to which it is assigned, and hence provides a measure of the quality of the clustering. The SSE is progressively reduced with each iteration.
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8.2.2.b.
Unsupervised clustering algorithm: ISODATA
ISODATA is an unsupervised clustering algorithm that aims to approximate the natural structure of a multidimensional dataset by iteratively passing through the data and defining clusters by minimizing the pixel separation values [Ball and Hall, 1965] (Figure 8-2). ISODATA is useful when little is known about the data prior to classification as it makes no assumptions on the underlying probability distribution of the datasets [Ball and Hall, 1967; Swain, 1972; Wu et al., 2008]. The user inputs the maximum number of clusters to identify (N), which allows the algorithm to choose N initial seeds that span the data space but need not correspond to real data values. These values will be equally separated along the vector that connects the minimum to the maximum data value in the multidimensional space [Huang, 2002].
In two dimensions, when the input
datasets are normalised, the N initial seeds will lie along the line joining (0,0) to (1,1). The position of the initial seeds is not crucial to the final clustering, provided that the algorithm is given enough processing time [Richards, 1995]. ISODATA adjusts the number of clusters iteratively and can merge and split clusters that have similar values [Ball and Hall, 1965; Palus and Bogdanski, 2003](similarity is described below) but will not output any more clusters than the maximum N. Additional input parameters are the convergence threshold- the maximum number of pixels which must be unchanging between iterations before the algorithm can cease- and the maximum number of iterations which the algorithm can execute [Ball and Hall, 1965; Richards, 1986; Swain, 1972]. The N initial seeds represent the centres of N initial classes. A pixel is then assigned to a class if its separation distance from the centre of the class is less than its distance to all other class centres. In two dimensions, the pixel separation from a class centre is given by [Richards, 1986]:
Equation 8-1:
ܦൌ ඥሺݔଵ െ ݏଵ ሻଶ ሺݔଶ െ ݏଶ ሻଶ
where D = separation distance; x = pixel vector (x1,x2); and s = cluster mean vector (s1,s2). Once pixels are assigned to one of the N classes, the class centres are recalculated by taking the mean of pixels assigned to each class in the previous step [Jensen, 2005]. The process then continues iteratively. The assignment to
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classes is independent of the contiguity of the pixels in the spatial frame and hence pixels within the same class can have a large spatial separation on the classified map. The pixels within each training cluster defined by ISODATA are used to calculate the mean and variance that class [Ball and Hall, 1965] which will provide input for the probabilistic classification algorithm MAXLIKE.
Figure 8-3: Schematic of the methodology employed in this study. The input data of thermal inertia and albedo, coupled with independent datasets, such as globally mapped properties of the surface and laboratory data, are used to interpret the classes. In this schematic most classes are composed of mixels and are therefore interpreted as a mixture of surface materials, with the exception of one (interpreted as pure dust).
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The input parameters used in this work to produce the thermophysical map of martian surface materials were: i.
Maximum of 10 classes
ii.
98% convergence threshold
iii.
Maximum of 200 iterations
The albedo and thermal inertia datasets were both normalized prior to classification by linear transformation to the range [0,1]. As albedo and thermal inertia have different physical meanings and different ranges of values across the martian surface, this normalization allowed them to contribute equally in the calculations of pixel separation [Richards, 1986]. The algorithm was not sensitive to the choice of (iii) as < 40 iterations were required to meet the converge threshold (Table 8-4) and thus the latter threshold determined the algorithm exit. The choice of maximum number of clusters is discussed below and in 8.2.3.a. The algorithm exited when at least 98 % of the pixels did not change between successive iterations. The classification is very sensitive to a higher convergence threshold. Table 8-4 shows the number of pixels classified differently when the convergence threshold is varied. Higher convergence threshold maps were not used for analysis for several reasons. Firstly, they allocated fewer of the 10 clusters to pixel values which occur in the low-mid latitudes, and instead partitioned the pixel values corresponding to the polar regions more finely. This occurred as the thermal inertia of the north polar region spans the full range of values. Increasing the number of clusters at the poles is not physically useful however as many of the north polar thermal inertia values are anomalous due to the data being taken near polar dawn when surface temperatures were changing rapidly. This resulted in the best fit thermal inertia from Equation 7-8 being unrepresentative [Putzig et al., 2005]. The polar regions of Mars are also less significant for habitability considerations due to their low temperatures, so a high level of detail on surface land cover is not required. Additionally, higher threshold maps appeared more pixelated at high latitudes indicating that the algorithm was unable to uncover coherent spatial patterns at those latitudes (again attributed to the quality of the polar data). The 98 % convergence threshold map provides the best quality map at non-polar latitudes, maintaining both the clear concentric patterns that are present in the lower threshold maps (see 8.2.3.d).
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Table 8-4: Sensitivity of classification to input parameters. Convergence threshold % 95 97.5 98 99 99.5
Iterations required to attain threshold < 10 < 10 107 pixels). Table 8-6: Approximating classification accuracy by χ2 with 2 degrees of freedom Confidence interval (%) 99.5 99 97.5 95 90 75 50 25 10 5
Chisquared value* 10.6 9.2 7.4 6.0 4.6 2.8 1.4 0.58 0.21 0.10
Pixel distance value 10.6 8.4 5.9 4.4 3.2 2.1 1.1 0.44 0.13 0.03
*Percentage of the χ2 distribution [Wackerly et al., 2002].
8.2.3.d.
Assessing classification accuracy
A statistical measure of confidence in the class assignment of each pixel is provided by the χ2 distribution with degrees of freedom (df) equal to the number of input parameters [Richards, 1986; Swain and Davis, 1978]. This provides an indicator of the fraction of pixels that are closer to the class centre and hence are more confidently classified. Pixels with smaller parameter distances (Equation 8-1) are more likely to be correctly classified as they lie closer to the mean of the Gaussian probability distribution function (PDF) of their class. Pixels that lie near the boundary of their class, and hence in the tail of their PDF, have a non-negligible probability of membership to an adjacent class. If the number of pixels in a given class is plotted against parameter distance from the class mean, the resulting histogram can be approximated by the χ2 distribution [Richards, 1986] (Figure 8-7;
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Remote Sensing of Thermal Materials on the Martian Surface
Table 8-6). A confidence interval of 95 % reflects that the probability of finding a pixel with a larger parameter distance and hence less confidently classified [Wackerly et al., 2002]. For example, the probability of finding a pixel with a parameter distance > 6 is ~ 5 %.
The χ2 statistics are generally applied to
independent variables, however thermal inertia and albedo are partially dependant.
Although the covariance of the (normalized) albedo and thermal
inertia datasets used in this study is non-zero (cov = 0.002), it is small compared to their individual variances, of 0.015 and 0.011 respectively. From Table 8-6 the distance values tend to fall below the χ2 distribution by < 30 % for confidence > 25 %, suggesting χ2 is a fair approximation. Hence the χ2 distribution can be applied to identify locations where the pixel assignment should be viewed with caution. This provides a measure only of relative classification confidence howeverwhether one pixel is more likely to be assigned to the correct class compared to another.
χ2 statistics do not provide a measure of absolute classification
confidence- whether the partitions between classes occur at meaningful values with the thresholds reflecting real trends in surface properties. This is discussed below.
Figure 8-7: Global distribution of pixel distance values (black solid line).
The
distribution is approximately χ2 as shown by the red hashed line (x-axis has been truncated as the real maximum parameter distance is 44.4).
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In Earth remote sensing, the accuracy of a classification is typically assessed through construction of an error matrix [Congalton, 1991]. The frequency of each class actually occurring on the ground (determined independently through field sampling of the study area) can be incorporated into the classifier as a weighted probability of class occurrence. Classification accuracy can be improved by the inclusion of these prior probabilities [Minggua et al., 2009; Strahler, 1980]- the likely coverage of each class on the ground. Prior probabilities assist in the separation of classes that share boundaries resulting in the boundary pixels having a similar probability of membership to both classes.
They are frequently
incorporated into a classifier such as Maximum Likelihood [Maselli et al., 1992; Strahler, 1980]. The strategy of including prior probabilities requires significant ancillary information on surface properties of the study area. On Mars the option of field sampling surface materials is unavailable (other than the 6 successful landing sites), nor are there any remotely sensed datasets which can provide an independent assessment of surface material grain size. The only method of truly assessing the accuracy of the thermophysical map produced here is through comparison to independent data sets on martian surface materials. Given a certain class in thermal inertia and albedo derived by the algorithmic classifiers, an interpretation is made of the dominant surface materials within that class using information from surface sampling and laboratory studies summarised in Table 7-3 (illustrated in Figure 8-3). To assess whether the physical interpretation of pixels within that class is accurate, independent data sets on surface materials within the spatial location of the class (8.2.1.b) must be analysed to validate the interpretation.
This methodology enables determination of whether the
thermophysical classes output by ISODATA and Maximum Likelihood are providing real information on surface materials. Validation of the thermophysical map is undertaken in 8.3.2. This ‘groundtruthing’ approach does not however provide an assessment of whether the division into 10 classes, shown in Figure 8-8, was an optimal partitioning of the data space. This is discussed in Appendix C.
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Figure 8-8: Scatterplot of all pixels showing the partitioning of the dataspace into 10 classes (thermal inertia x-axis is linear above; logarithmic below). The discrete values of thermal inertia below 50 tiu are an artefact of the ‘lookup’ model used to derive thermal inertia (Equation 7-8).
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8.3.
Results: the spatial distribution of surface materials
The 10 thermophysical units mapped on the martian surface are shown in Figure 8-9 with their corresponding attributes given in Table 8-7. The classification confidence of the thermophysical map is shown in Figure 8-10, where the pixel colour reflects the classification confidence given in Table 8-6.
The poorest
classified 1 % of pixels have a widespread class membership with no strong relationship to any class (class 5: 11.3%, class 3: 13.6%, class 10: 14.8%, class 1: 18.3%, class 9: 21.2% and less than 8% membership in each of the remaining classes). They occur in the regions of Valles Marineris, Aram Chaos, eastern Acidalia Planitia, Hellas and Argyre, Olympia Planum and Syrtis Major.
The
difficulty in the algorithmic assignment of these pixels to classes indicates that the pixel values occurred at the boundary of, and hence had a comparable probability of membership to, more than one class. Given that the pixels with the lowest classification confidence generally have wide spatial separation however, they will not have a substantial effect on the interpretation of surface materials given here, with the exception of the expanse in Syrtis Major. Table 8-7: Attributes of the 10 thermophysical classes (in the form: mean, standard deviation; minimum-maximum). Class 1 2 3 4 5 6 7 8 9 10*
Thermal inertia (tiu) 269, 79 5-875 251, 54 33-480 233, 75 5-648 162, 76 5-431 197, 85 5-584 214, 81 5-761 130, 72 5-470 67, 40 5-364 392, 313 4-1580 3620, 989 37-4999
*Ignoring pixels forming border
Albedo 0.12, 0.01 0.07-0.16 0.15, 0.01 0.14-0.16 0.16, 0.01 0.15-0.19 0.18, 0.01 0.17-0.20 0.21, 0.01 0.19-0.22 0.23, 0.01 0.21-0.26 0.26, 0.01 0.24-0.28 0.29, 0.01 0.27-0.32 0.35, 0.06 0.19-0.57 0.12, 0.15 0.05-0.53
Elevation (km) -0.8, 3.1 -6.7-6.9 -0.1, 2.6 -7.0-7.6 -0.4, 2.7 -8.1-8.2 0.1, 2.6 -8.2-8.8 -1.5, 3.1 -8.1-20.9 -1.9, 2.7 -7.8-21.1 -1.4, 2.9 -6.6-21.2 0.4, 3.1 -6.5-21.2 -1.2, 3.5 -7.8-18.0 -2.1, 2.7 -8.1-6.3
Dust cover index 0.973, 0.006 0.852-0.991 0.971, 0.006 0.851-0.991 0.968, 0.007 0.851-0.991 0.965, 0.008 0.851-0.991 0.959, 0.009 0.851-0.991 0.948, 0.012 0.851-0.991 0.940, 0.011 0.851-0.991 0.935, 0.009 0.851-0.991 0.941, 0.011 0.880-0.991 0.965, 0.009 0.856-0.991
% Map (pixels) 11.2 12.4 12.8 11.3 10.7 14.2 8.8 10.1 3.1 2.2
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Remote Sensing of Thermal Materials on the Martian Surface
On the next pages: Figure 8-9: Map of the martian surface classified into 10 thermophysical units. Significant surface regions and landing sites are named. The map reveals clear global spatial patterning in the identified classes. The corresponding classification confidence of each pixel is mapped in Figure 8-10. Figure 8-10: Classification confidence thresholds based on pixel parameter distances. Bin colours reflect the percentage of pixels that are classified more accurately. For example, the 99 % bin reflects the poorest 1 % of pixels. Syrtis Major has the largest expanse of poorly classified pixels, and hence should be interpreted with caution.
Figure 8-9
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Remote Sensing of Thermal Materials on the Martian Surface
Figure 8-10
240
241
Table 8-8: Interpretation of the 10 martian thermophysical classes. Class 1
Spatial description Bordered by 2; Acidalia Planitia, Olympia Planum, Thaumasia Planum, Syrtis Major, Tyrrhena Terra
2
Bordered by 1 and 3; Chryse Planitia, Terra Sirenum
3
Bordered by 2 and 4; Argyre
4
Bordered by 3 and 5; Promethei Terra, Arcadia Planitia
5
Bordered by 4 and 6; Hellas Planitia; Marikh Vallis
6
Bordered by 7 and 5; Xanthe Terra, Isidis Planitia, Amenthes Planum Bordered by 6 and 8; Aolis Planum, Terra Sabaea, Amazonis Planitia Bordered by 7; Tharsis, Olympus Mons, Elysium, Arabia Terra Polar regions
7
8
9 10
Appears sparsely in northern polar terrain, Valles Marineris, Hellas & Argyre; also includes map boundary (albedo = 0; thermal inertia = no data)
* See section 7.1 for the basis of this interpretation.
General interpretation* Sand, duricrust and rocks. Sand predominantly around 400 µm grains; includes a small fraction of coarse sand/pebbles/duricrust up to 3 mm grains; also consistent with ice cemented soil under ~ 2 µm of dark fines; minimal dust Sand. Similar to class 1 but distinguished by thin and sporadic covering of dust deposits 2 cm thick in some regions; mostly fine sand < 500 µm; some coarse sand/pebbles up to 2 mm grains Dust. A > 2 cm thick layer.
Ice, dust, some sand. Some sand, rocks and bedrock < 2 m across under bright dust or frost Ice and rocks. Dark clean rocks under thin coating of dust or frost several microns thick. Also includes error and boundary pixels.
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Remote Sensing of Thermal Materials on the Martian Surface
Figure 8-11: Histograms of martian thermophysical classes 1-8 in thermal inertia (left) and albedo (right) data. Colours as in Figure 8-9. The location of the peaks allows the relative fraction of different thermal surface materials to be distinguished.
243
Figure 8-12: Histograms of martian thermophysical classes 9 & 10 in thermal inertia (left) and albedo (right) data. The colour coding matches Figure 8-9.
8.3.1. Global trends The thermophysical map reveals a pattern of concentric envelopes in the spatial occurrence of the classes in both hemispheres. The histograms of albedo and thermal inertia values within each class in Figure 8-11 and Figure 8-12 reveal that this sequence corresponds to a decrease in albedo moving outwards from the equator through the sequence of classes 8 ื 7 ื 6 ื 5 ื 4 ื 3 ื 2 ื 1. The mean thermal inertia also generally increases through this sequence (with the
exception of classes 4 and 5 which have lower values, see Table 8-8). This global trend supports the expected physical interpretation of thermal inertia-albedo units as reflecting variations in grain size and grain cementation (8.1). The trend of fine dust-size grains being dominant in the low latitudes, coarse sand in the midlatitudes, and ice at the high latitudes, is consistent with previous thermophysical unit maps [Mellon et al., 2000]. Furthermore, the boundaries of the classes in
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Table 8-8 are broadly consistent with the decreasing coverage of surface dust with distance from the equator, evidenced by the increasing value of the mean dust cover index through the sequence of classes 8 ื 1 [Ruff and Christensen, 2002].
Hence the thermophysical map displays the expected global trends in surface grain size. Larger particles such as rocks and bedrock are largely unimportant in the interpretation of the classes as they are either a small fraction of each pixel or are blanketed by centimetres of fines (discussed in 7.1.2).
8.3.1.a.
Comparison between thermophysical classes & major geological
structures The delineation of major geologic structures such as Valles Marineris, Olympus Mons, and a number of large impact craters in the thermophysical map suggests a broad correlation between the classes and martian geology. Some geologic units have distinct thermal behaviour from their surrounding terrain and hence are distinguished in the map. The Valles Marineris region is shown in Figure 8-13 with the major canyons being outlined by a single geologic unit (purple) and predominantly in-filled by two distinct geologic units (pale yellow and blue). In the thermophysical map, the major canyons and the western labyrinth of valleys are clearly defined by a boundary of class 1 (red) and class 10 (black). The canyons are also in-filled by two classes of surface materials – class 2 and 3. Furthermore, the boundary between classes 4 and 5 (light and dark green) adjacent to the channels is located near the boundary between the low viscosity lava flows of the ‘ridged plains unit’ and the volcanic flows of the ‘syria planum formation’ in the geologic map [Nimmo and Tanaka, 2005; Scott et al., 1986; Skinner et al., 2006]. This provides some validity to the thermophysical map and suggests that it may be used to resolve different types of lava flows, however the Valles Marineris region of the map has a lower classification confidence in Figure 8-10 and so should be viewed with caution.
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Figure 8-13: Valles Marineris. A comparison of the canyon system in the thermophysical classification map (top), geologic map (middle), and the geologic map classified into Amazonian (blue), Hesperian (red) and Noachian (green) aged terrain [Nimmo and Tanaka, 2005; Skinner et al., 2006]. Noctis Labyrinthus, Ius Chasma, Candor Chasma, Coprates Chasma and Melas Chasma are clearly delineated in both the geologic and thermophysical maps. Scale bar is given on the top map.
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Remote Sensing of Thermal Materials on the Martian Surface
Figure 8-14: The summit of Olympus Mons is distinguished in the thermophysical map. Olympus Mons can also be identified by concentric distributions of fine grained material (classes 6, 7 and 8) at its summit (Figure 8-14). A number of impact craters with diameter over 50 km are distinguished in the thermophysical map, including: Micoud, Lyot, Moreux, Nier, Mie, Stokes, Milankovic, Bernard, Escorial, Timoshenko, Korolev and Lomonosov.
These craters can be identified by
concentric circular structures of thermophysical units that contrast with the units dominating the surrounding terrain.
This is consistent with observations of
distinct high thermal inertia impact ejecta surrounding many martian craters (8.1.1). The latter two craters are discussed in more detail below. Little correlation with elevation is seen in the spatial locations of the 10 classes as all classes occur in both hemispheres and across a wide range of surface elevations (Table 8-8), consistent with previous derivations of thermophysical units (8.1.1). However, specific elevation domains can be recognised, with the Tharsis Montes, Elysium Mons, Alba Mons and Ascraeus Mons all falling within class 8, possibly because they are thickly covered with dust [Basilevsky et al., 2005].
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8.3.1.b.
Comparison between thermophysical classes & geologic age
Martian terrain is categorised into three broad periods of geologic history based on impact crater densities. Surface crater counts reflect the age of the surface since its last significant reworking [Barlow, 2008], with older surfaces maintaining a more complete cratering history and hence having high crater numbers. The age bands are Noachian (surface ages 4.1-3.7 Ga), Hesperian (3.7-3.0 Ga) and Amazonian (3.0-0 Ga) [Hartmann and Neukum, 2001]. Each geologic period is characterised by different surface processes and hence a weak relationship between surface age and grain size may be expected. Noachian surfaces, being the oldest, are heavily cratered and degraded. During this period the surface experienced extensive liquid water erosion through major flooding, such as the events that carved Valles Marineris [Dohm and Tanaka, 1999] and other valley networks [Fassett and Head, 2008a, 2008b], and likely had long-term standing water to produce the observed sedimentary layers [Malin and Edgett, 2000] and clay minerals [Bibring et al., 2006]. Hesperian surfaces also experienced water activity, with outbursts of water erosion forming the outflow channels [Fassett and Head, 2011] and acidic waterrock interaction leading to the sulfate mineralogy [Bibring et al., 2006]. Volcanic activity was frequent during this period, with extensive lava plains covering the surface [Carr and Head, 2010].
The Amazonian epoch is characterized by
significantly less water and lava erosion, although some examples exist [Fuller, 2002],
with
predominantly
water
poor
glacial/periglacial activity [Head et al., 2006].
environments
but
extensive
The primary weathering was
chemical with the oxidation of iron bearing surface rocks [Bibring et al., 2006]. Figure 8-13 suggests that that the thermophysical units may be sensitive to geologic terrain age, as the borders between Noachian, Hesperian and Amazonian units are closely matched by class boundaries. To further explore the potential of the thermophysical map to identify surface ages, the division of terrain from each of the three geologic age groups into the 10 classes was examined. The geologic map listed as dataset 7 in Table 8-3 contained 2878 mapped polygons of which 2202 had designated terrain ages from crater counts (marked by ‘A...’, ‘H...’ or ‘N...’ in ‘Unit Symbol’).
Figure 8-15 reveals that although each class is comprised of
terrain of all surface ages, there is a clear mapping from surface age to
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thermophysical unit. No previous mappings of thermophysical classes reported similar trends.
Noachian terrain occurs predominantly in classes 2 and 3;
Hesperian in class 6 and Amazonian in class 8. The latter relationship is due in part to the thick mantling of dust characterising class 8 (Table 8-8), which obscures part of the cratering record of the surface [Hartmann and Neukum, 2001]. The former relationships will be briefly discussed below.
Figure 8-15: Relationship between terrain age and thermophysical units. For each geologic epoch percentages are calculated as the total surface area of terrain that falls within each class, out of the total global area of terrain of that epoch.
8.3.2. Interpretation of thermophysical classes The mixture of surface materials within each thermophysical unit is interpreted in Table 8-8 by comparing the modes of thermal inertia and albedo within the class (Figure 8-11, Figure 8-12 and Table 8-8) to end-member martian materials in Table 7-3. Approximately ~ 35 % of pixels in the global map are dominated by a single type of surface material and hence can be unambiguously interpreted (Table 8-2). The remaining 65 % of the map are mixels, containing a mixture of different materials within the spatial resolution of the pixel. Hence most classes in Table 8-8 are interpreted as a mixture of more than one surface component. Within each class, the horizontal and vertical (within the sensing depth) spatial coverage of different surface materials determines the location and strength of the peaks in
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albedo and thermal inertia.
This allows the fraction of each material to be
constrained. Duricrust cannot be unequivocally identified from thermal inertia and albedo values as they do not allow the difference between large grains and small grains cemented together to be distinguished. Below, the validity of the thermophysical map is examined by comparing the interpretation of each class with independent observations. The classes are examined in numerical order.
8.3.2.a.
Classes 1, 2 & 3: sand
Class 1 includes the majority of pixels with extremely low albedo (< 9 %) which form the tail of the global albedo histogram (Figure 7-2). From the thermal inertia values within the class, grain diameters are predominantly less than 1 mm and hence the surface is dominated by sand. The class also includes a small fraction of very coarse > 4 mm sand grains or duricrust resulting in the high thermal inertia values > 350 tiu (an alternative explanation is provided in 8.3.2.b).
A large
component of rocks and bedrock within each pixel can be excluded as there are no values above 900 tiu [Mellon et al., 2009b]. The low albedo excludes dust [Jakosky, 1986]. Class 1 occurs over an extensive range of latitudes including the regions of Syrtis Major, Acidalia, Solis Planum and Tyrrhena Terra and infilling Valles Marineris. Although the last example occurs on Amazonian aged surfaces, the majority of class 1 terrain occurs on Noachian and Hesperian aged terrain (Figure 8-15). Class 2 has a similar interpretation with three distinctions: (i) a smaller fraction of large grained > 400 tiu materials; (ii) a slightly higher fraction of bright dust evidenced by the higher albedo mode; and (iii) class 2 occurs more frequently on older Noachian terrain. Class 3 is similar to class 2 but has a lower fraction of medium-coarse sand and a higher fraction of fine sand and dust grains < 100 µm across. If the interpretation of classes 1-3 as being sand dominated is correct, then the location of the classes should be correlated with expanses of sand visually identified from imagery. Dunes larger than 1 km2 have been mapped from THEMIS and MOC imagery [Hayward et al., 2008] and the dune boundaries can be intersected with the thermophysical map to measure the overlap with different classes. From the ~ 105 km2 area of mapped dune coverage, encompassing 547 distinct dune fields [Hayward et al., 2007, 2008], between ± 65˚, 95 % of total dune
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area was found to occur in classes 1-3, representing 1.6×105 km2 (Figure 8-16). This strong correlation between large dune occurrence and class 1-3 terrain provides strong supporting evidence for these surfaces being sand dominated. In comparison, classes 4-10 have few medium-large dunes. An example of a dune field in class 1 and 2 terrain is the Kaiser dune field [Edgett and Christensen, 1991] shown in Figure 8-17.
Figure 8-16: Fraction of global dune area included within each cluster. The absence of mapped dunes within a class does not imply that dunes do not exist as they may be smaller than the 1 km2 resolution of imagery used to construct the database. Furthermore, surfaces can have a high sand fraction without saltation driving the formation of dunes. Class 2 and 3 terrain was sampled at Meridiani Planum by the Opportunity Rover (Figure 8-18), which found the surface to be predominantly dust free with an albedo of 0.12 [Fergason et al., 2006; Glotch and Bandfield, 2006]. The terrain at Meridiani is dominated by basaltic sand and grey spherical hematite grains, millimetres in diameter [Ruff et al., 2008]. The high thermal inertia materials observed in the landscape were sparse rocks (400-1100 tiu) and duricrust [Golombek et al., 2008b; Mellon et al., 2008b].
This is consistent with the
interpretation given in Table 8-8, providing further validation of the dominant surface materials within the classes 2 and 3.
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In at least one region, the sand-dominated class 1 terrain obscures a deeper water ice-cemented layer. The Olympia Undae region within Olympia Planum at ~ 82 ˚N has a large occurrence of class 1 terrain (classes 2, 3 and 10 are also present with small spatial coverage) and is characterised by a vast sand sea [Byrne and Murray, 2002; Langevin et al., 2005a] surrounding the perennial water ice polar cap. Hydrogen mapping of these dunes indicates that although desiccated in the top ~6 cm and hence throughout the sensing depth of TES in sand, the dunes harbour ~ 30 wt% of hydrogen beneath [Feldman et al., 2008]. This reinforces the limitations of extrapolating the surface material interpretations given here to depths beyond the diurnal skin depth.
Figure 8-17: Kaiser dune field. The mapped boundary of the dune field is shown in grey, and overlaps class 1 and 2 terrain. The dune field is visible in infrared THEMIS images18 and in the higher resolution HiRISE red images overlaid19 with the black rectangle giving the extent of the HiRISE inset20.
18
THEMIS images: I00905006, I00880002; metadata in Appendix A, Table A2; processing details in Appendix
B. HiRISE images: ESP_013017_1325, ESP_015918_1325; metadata in Appendix A, Table A1; processing details in Appendix B. 20 HiRISE image ESP_011883_1325; metadata in Appendix A, Table A1; processing details in Appendix B. 19
Remote Sensing of Thermal Materials on the Martian Surface
Figure 8-18
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On the previous page: Figure 8-18: Classification of the six martian landing sites. Only Opportunity is currently active on the surface and journeying towards Endeavour crater.
8.3.2.b.
Classes 1, 3 & 6: pebbles, duricrust.
In Table 8-2 the thresholds of thermal inertia and albedo that would lead to an interpretation of the surface as duricrust/pebbles/coarse sand were given and the occurrence of these values was mapped, revealing that the values occur exclusively in class 1.
The thermal inertia values between 500-600 tiu are likely to be
duricrust or coarse sand, whereas the higher values up to 800 tiu are likely to indicate pebbles/rocks (Table 7-3). Although the spatial coverage of these values within the class is extremely limited- with the most extensive regions occurring in Valles Marineris and Acidalia Planitia- there are a number of significant correlations with surface features. Two examples are given in Figure 8-19, where the occurrence of high thermal inertia values in class 1 are contiguous with distinct surface morphologies: (i) infilling a Hesperian aged outflow channel in Aram Chaos [Skinner et al., 2006] and (ii) associated with large blocky material within an unnamed crater in northern Syrtis Major. Additionally, class 1 dominates Nili Patera (within Syrtis Major) which is known of have vast exposures of rock [Christensen et al., 2003], suggesting that the interpretation of this unit as rocky (pebbles) may be more valid than coarse sand/duricrust. Regardless, this is a significant result as the thermophysical map is providing a higher level of detail on coarse surface materials than previous maps (8.1).
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Figure 8-19: Duricrust or pebbles? THEMIS infrared images21 are overlain on the geologic map. The red pixels are the high thermal inertia values within class 1 discussed in the text. Duricrust/pebbles/coarse sand are associated with ‘warm’ blocky regolith in the left, and channel material in the right. From Table 7-3, duricrust is most likely to have thermal inertias in the range of 350-600 tiu, although these values can also reflect the average of sub-pixel mixtures of pebbles and sand. The presence of duricrust on the surface therefore cannot be unequivocally determined solely through albedo and thermal inertia, and there may be duricrust within each of the classes. Furthermore, duricrust cannot be excluded as an explanation for thermal inertia values < 350 tiu, as it may THEMIS images: I09042008 (left); I06461011, I00931002 (right); metadata in Appendix A, Table A2; processing details in Appendix B. 21
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be obscured by several skin depths of sand or dust. The remote detection of localised expanses of duricrust and pebbles is significant, as generally the fine component of the surface obscures larger grains (7.1.4). From Table 8-8 classes 1, 3 and 6 have orbital values of thermal inertia that are consistent with duricrust and/or pebbles (thermal inertias > 600 tiu Table 8-7, grains > 4 mm), mixed with some fraction of dust (depending on the value of albedo). Without groundtruthing duricrust, pebbles and mixtures of rocks with coarse sand, all remain plausible explanations. The distinction is important, as it effects the roughness and hence navigability of the surface for landing missions [Steltzner et al., 2010]. In at least one location the aforementioned occurrence of duricrust and rocks has been observed. The Spirit landing site within class 6 terrain (Figure 8-18) was dominated by a < 1 mm thick bright dust covering [Fergason et al., 2006] over pebble-rich terrain and drift deposits (particles < 100 µm) [Golombek et al., 2006, 2008b]. This is consistent with the surface material description of class 6, and validates the interpretation given in Table 8-8. Spirit also observed a strong presence of duricrust (200-300 µm cemented grains) [Golombek et al., 2006, 2008b]. The duricrust was indistinguishable however in orbital thermal inertia (< 430 tiu [Christensen et al., 2004a]) and albedo (0.22 [Arvidson et al., 2004]) from bright dust coating coarse sand or pebbles. In other regions of class 6 however, the pebble component can be distinguished in the orbital thermal inertia as discussed above. Opportunity will be able to observe the dominant surface grain size and the presence of duricrust at its destination of Endeavour crater [Arvidson et al., 2011] within class 2 (Figure 8-20). Sand in class 2 will also be sampled as dunes have been imaged within Endeavour [Bridges et al., 2011].
8.3.2.c.
Classes 4-7: fine-medium sand & pebbles, varying dust
Classes 4-7 span the region of intermediate thermal inertia and albedo values that can be due to varying mixtures of a number of surface materials. Each class is interpreted as surfaces composed of a mixture of fine-medium sand, dust and some pebbles, but the classes differ in the relative fraction of each component, which is deciphered in Table 8-8 from the location of the modes in thermal inertia and albedo. For example, class 4 is dominated by mixtures of fine sand and dust with dominant grains sizes around 10 µm (thermal inertia mode at 80 tiu) and 150 µm
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(mode at ~ 190 tiu). As clean martian sand is dark, the albedo > 0.15 of class 4 suggests the ubiquitous presence of bright dust microns thick- brightening the albedo and slightly lowering the averaged thermal inertia across the pixels. Similar logic can be applied to classes 5-7, with the results summarised in Table 8-8. Although encompassed by only a small fraction of pixels, the presence of coarse sand, pebbles, and possible duricrust within classes 4-7 is indicated by thermal inertia values > 350 tiu. The fraction of these coarse-grained or cemented materials decreases towards the equator (with increasing class number). Several of these classes have been sampled by landers on the martian surface. Class 6 was discussed above (8.3.2.b). Pathfinder sampled near the border of classes 3 and 4 (the low resolution of the thermophysical map prevents a more precise determination) in Ares Vallis, finding the surface to be dominated by finegrained drift material and sand [Matijevic et al., 1997], with ~ 16 % of the observed area containing semi-rounded pebbles and larger rocks [Golombek, 1997]. Dark rocks were observed to have discontinuous coatings of bright red dust, raising their albedo [Smith et al., 1997]. The interpretation of grain sizes within class 3-4 terrain is in agreement with the fine component observed on the surface [Bell et al., 2000] as this dominates the orbital thermal inertia [Putzig and Mellon, 2007a]. As the Pathfinder site had the highest rock abundance of all of the landing sites [Golombek et al., 2003], but there are no pixels within class 3 or 4 that have an orbital thermal inertia consistent with pebbles or larger rocks, it is not surprising that there are very few pixels globally that can be uniquely determined as rock (Table 8-2).
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Figure 8-20: Path of the Opportunity (above) and Spirit (below) rovers. Spirit has come to rest at Home Plate.
Opportunity ventures towards Endeavour crater,
transitioning from class 3 to class 1 terrain. Mars Digital Image Map 2.1 in Google Mars image with .kmz waypoints superimposed [Archinal et al., 2003].
8.3.2.d.
Class 5: possible subsurface ice
Although the global interpretation of materials within class 5 was presented above, groundtruthing suggests an alternate explanation for at least some occurrences of class 5 terrain. This class was sampled by the Phoenix lander at 68˚ N (Figure 8-18). In orbital thermal inertia and albedo class 5 is dominated by a ubiquitous coating of dust grains < 10 µm across, over sand grains < 1-3 mm. Beneath these
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fines however, the Phoenix Lander observed solid water ice exposed from under ~ 5 cm of dry desiccated material by the landing thrusters [Smith et al., 2009b]. Furthermore, the soil was observed to become transiently ice-cemented [Arvidson et al., 2009] on diurnal time scales [Mellon et al., 2009a] with a stable ice layer occurring 10 cm beneath the surface [Sizemore et al., 2010]. Similar ice-table depths would be expected in other locations [Mellon et al., 2008a] at these latitudes until winter allows the seasonal stability of ice on the surface [Cull et al., 2010; Searls et al., 2010]. Given that the diurnal skin depth in desiccated fines < 1 mm across is < 6 cm (Table 7-2), it is likely that the orbital thermal inertia values with class 5 contain a contribution from the ice-cemented soil. This raises the question of whether other locations of high-latitude class 5 terrain, such as Lomonosov crater (Figure 8-21), are also associated with subsurface ice. Although addressing this point thoroughly is beyond the scope of this work, a few features are worth noting. Figure 8-21 reveals an occurrence of class 5 inside the large impact crater Lomonosov which is potentially correlated with water ice. The presence of ice-cemented soil inside Lomonosov is speculated from thermal observations [Christensen et al., 2001a]. Water frost is certainly present seasonally in the interior of the crater (http://themis.asu.edu/zoom-20040616A). Other lower latitude occurrences of class 5 include east of Elysium Mons, Marikh Vallis, Solis Planum and Hellas Planitia, some of which show morphological indicators of past water ice [Costard, 1989; Depablo and Komatsu, 2009; Dundas et al., 2007; Lucchitta et al., 1986; Soare et al., 2007].
There is no current evidence for the present-day existence of
subsurface ice at these locations due to the predominantly low water-equivalent hydrogen of 2-4 wt% (with the exception of Marikh Vallis which shows an elevated abundance of 6 wt%; [Maurice et al., 2011]). Nevertheless it would be worth examining in future studies whether there is any spectral evidence for water ice in shallow soil within low-latitude occurrences of class 5.
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Figure 8-21: Lomonosov crater in the thermophysical map (above) and geologic map (below) with THEMIS infrared images superposed22. The geologic units are ‘impact crater material, superpose’ in yellow (age unconstrained) and ‘Vastitas Borealis Formation, knobby member’ in green (Hesperian). The red rectangle indicates the extent of the upper map. Diameter of Lomonosov crater is ~ 150 km.
8.3.2.e.
Class 8: dust
The thermal inertia and albedo values within class 8 are unambiguously consistent with bright dust grains < 10 µm across, forming a surface layer > 2 cm thick (diurnal skin-depth in dust, Table 7-2) and hence dominating the TES thermal inertia measurements. This interpretation is validated by the low mean DCI of the class [Ruff and Christensen, 2002] indicating that dust dominates the spatial coverage of the pixel. Dust is clearly distinguished on Mars from orbital thermal inertia values around 55 tiu and an albedo around 0.27 (Table 7-3). A strong peak THEMIS images: I04919006, I04507005, I04482005, I04407005, I03658006; metadata in Appendix A, Table A2; processing details in Appendix B. 22
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Remote Sensing of Thermal Materials on the Martian Surface
at the latter value is clearly seen in the global distribution of the albedo (Figure 7-2) as dust is a major global surface component.
Hence it is an important
accuracy test for the classification algorithms that this material is clearly distinguished in known regions by one of the thermophysical units. Radar, visual and infrared observations have identified metres thick accumulations of dust on the martian surface in regions such as Amazonis Planitia [Christensen and Moore, 1992] and Arabia Terra [Mangold et al., 2009] - both contained within class 8 and hence providing further validation for the interpretation of the dominant surface materials within this class. Class 8 occurs at low latitudes around the young Amazonian volcanic structures of Olympus Mons, Tharsis Montes and Elysium Mons– consistent with independent observations of thick dust in these regions [Christensen, 1986a].
8.3.2.f.
Classes 9 & 10: ice, polar caps & bedrock
The thermophysical map in Figure 8-9 provides remote detection of surfaces dominated by water ice- classes 9 and 10 (the ‘lookup model’ used to derive thermal inertia excludes surfaces cold enough to have CO2 ice; 8.1). Thermal inertia values > 1000 tiu with corresponding high albedos > 0.27 are indicative of either rock covered by a thin coating of dust or solid ice. These values occur exclusively in classes 9 and 10 and are predominantly located at > 75˚ latitude in both hemispheres. The spatial locations of classes 9 and 10 are consistent with: (a) the distribution of subsurface water concentrations > 50 wt% [Boynton et al., 2002b; Mitrofanov et al., 2002] (see 5.2); (b) the observations of shallow solid water ice in the northern polar terrain [Bandfield, 2007]; and, (c) the extent of the polar caps
[Titus et al., 2003; Zuber et al., 1998].
The correlations with
independent observations of the surface validate the interpretation of classes 9 and 10 as dominated by H2O ice. A specific example is given in Figure 8-22. The interior of Korolev crater shows ice-related morphologies on the interior mound [Burr et al., 2009] (pingoes; 5.2.4.b) that have been found from spectra to have a water ice composition [Brown et al., 2007]. From Figure 8-22 class 9 in-fills Korolev crater, and hence its occurrence is closely associated with the exposed ice [Armstrong et al., 2005] providing further validity to the interpretation.
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Some low latitude occurrences of class 9 (for example around 8˚N, 251˚E) could potentially indicate low-latitude ice cemented terrain remaining from periods of low obliquity [Levrard et al., 2004]. One example occurs in the ‘Olympus Mons Formation’ [Skinner et al., 2006] shown in Figure 8-14.
All low latitude
occurrences of class 9 however have low thermal inertia < 150 tiu but high albedo > 0.3, which is more likely to be consistent with dust with an elevated albedo due to processing errors or specular reflection.
Figure 8-22: Korolev crater in the thermophysical map (above) and geologic map (below) with THEMIS infrared image overlaid23. Geologic units are ‘impact crater material, superpose’ in the yellow and ‘mantle material’ (Amazonian) in the teal. The red rectangle indicates the extent of the upper map. Crater diameter is ~ 84 km.
23
THEMIS image I04526006; metadata in Appendix A, Table A2; processing details in Appendix B.
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The thresholds given in Table 8-2 to encompass rock occur within class 10, as this class includes the newly derived extremely high thermal inertia values [Putzig and Mellon, 2007a].
If these values of 4000 tiu and above are real, then by
extrapolating from Equation 7-9 they reflect pixels dominated by rocks larger than > 0.6m [Golombek et al., 2003]. Although some of those pixels occur in the Valles Marineris canyon system (Figure 8-13), the vast majority occur in the polar terrain above 80˚N. This is consistent with observations of martian boulder fields, such as in the northern plains [Golombek et al., 2008a; Seelos et al., 2008]. Some of the extensive exposures of these high thermal inertia pixels may however be representative of the vertical layering of ice beneath dark sand deposits which is a common stratigraphy at high latitudes [Byrne and Murray, 2002].
8.3.3. Comparison to previous work To highlight the new information provided by this thermophysical map, a comparison will be briefly made to the most recent previous thermophysical map [Putzig et al., 2005] (Figure 8-1). Given the differences between the methodology employed in this work and that of Putzig et al. [2005], the regions of expected agreement between the two maps can be inferred.
Previous studies chose
thresholds for the thermophysical classes by isolating the modes in the global histograms of martian thermal inertia and albedo (Figure 7-2). The boundaries of those units were refined by observing the density of points in a contour plot. Close agreement between both Putzig et al. [2005] and this work is therefore expected in the identification of dust (class 8 in Figure 8-9), sand (classes 1-3) and ice (class 9) dominated terrain (8.1.2). There is one caveat to the expected regions of similarity between this work and that of Putzig et al. [2005]. This study utilised the updated thermal inertia dataset derived from a more accurate lookup model [Putzig and Mellon, 2007b].
The 2007 model allowed values as large as 5000 tiu to be
calculated, whereas the maximum in the earlier model was 800 tiu [Putzig, 2006]. Thermal inertia values over 800 tiu occur in classes 9 and 10, in the polar region and Valles Marineris, and hence a significant difference between the two studies in the classification of pixels in these regions is expected. The 2007 thermal inertia map also contained a much higher fraction of real data (8.2.1.a).
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Table 8-9:Comparing the partitioning of pixels into classes in this study and in Putzig et al. [2005]. Putzig % Membership in each cluster Class 1 2 3 4 5 6 A B C D E F G N/A % of total map
% of total 7 8 9 10 map 0 0 0 0 0 2.1 29.6 68.1 0.2 0 12.9 24.0 32.9 29.3 12.0 1.7 0.1 0 0 0 0 33.0 0 0 0.3 2.1 23.9 56.7 16.5 0.1 0.4 0 19.0 1.2 2.4 12.7 60.0 21.7 1.1 0.3 0.2 0.4 0 4.1 96.1 2.8 0.4 0 0.2 0.2 0.3 0 0 0 0.3 20.6 10.0 12.2 5.5 16.6 17.6 4.8 1.6 9.8 1.3 6.1 0.1 0 0.2 0.2 1.4 41.6 14.6 29.2 12.7 0 0.6 6.9 3.5 7.4 17.3 15.4 7.4 6.2 4.4 9.6 21.9 24.1 11.2 12.4 12.8 11.3 10.7 14.2 8.8 10.1 3.1 5.4 N/A
*Total number of pixels in both maps is 2.592× 107. 24.1% of pixels in the map were not partitioned into classes by Putzig et al. [2005].
The closest agreement between this work and that of Putzig et al. [2005] relates to class E and class 1 (96.1 % agreement; Table 8-9). Class 1 includes the majority of pixels with extremely low albedo (< 9 %) that form the tail of the global albedo histogram and hence are easily distinguishable through the methods of both studies. Putzig et al. [2005] interpreted this unit as sand, rocks and bedrock with little or no fines. From the thermal inertia values within class 1, grain diameters are predominantly less than 4 mm and hence the surface is dominated by sand with a small fraction of very coarse > 1 mm sand grains or duricrust or pebbles. The high degree of overlap between class E and class 1 therefore indicates that Putzig et al.’s [2005] class E was dominated by sand-sized particles. The next closest agreement occurs between class A and class 8 (68.1 % agreement). Putzig et al. [2005] interpreted this unit as ‘bright unconsolidated fines’, which aligns with the interpretation of class 8 being bright dust grains < 10 µm across forming a layer > 2 cm thick. Dust is clearly distinguished on Mars remotely by a strong peak at albedo 0.27 and thermal inertia 55 tiu. Thus the occurrence of thick dust deposits was easily distinguishable through the methods of both works. Putzig et al.’s [2005] class A also included pixels with lower albedo and higher thermal inertia, indicative of fine sand particles up to ~ 60 µm across.
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In this work these pixels have been incorporated into a distinct surface unit, class 7, with a few μm-thick cover of bright dust over darker fine to coarse sand. The good classification agreement for sand (class 1) and dust (class 8) with Putzig et al.’s [2005] classes is as expected. A similar level of agreement would be expected for ice (class 9) but Putzig et al. [2005] did not define a separate class for ice. The vast majority of pixels in the global map will be mixels, with their thermal inertia and albedo corresponding to an average of more than one component surface material.
For this reason, algorithmic classification was applied to
determine class boundaries. This technique groups similar pixels together by minimising the spread in albedo-thermal inertia space within each class. Hence in some regions the classification provided a significantly different partitioning of thermal inertia and albedo, summarised in Table 8-9. There is moderate agreement between Putzig et al. [2005] class B and classes 1-3 of this study (~ 88% of class B in these classes). Putzig et al. [2005] interpreted class B as sand, rocks, bedrock, some duricrust. Class 1-3 were shown to be surface dominated by fine to coarse sand with varying fractions of dust, due to their strong correlation with martian sand dunes. Hence class B is dominated by sand. The high thermal inertia values within class B are incorporated into class 1, and were interpreted as indicative of duricrust/coarse sand/pebbles < 3 mm across. Class D of Putzig et al. [2005] was described as dark regolith with a low density from the loss of subsurface ice. Here, class D is predominantly partitioned into class 4, which is consistent with mixtures of sand and dust.
Although
duricrust cannot be eliminated from any of the classes in this study if obscured by fines, classes 1 and 6 showed surfaces which could be identified as clean duricrust. Alternate surface materials such as coarse sand, pebbles, or ice-cemented soil under dark fines cannot be ruled out. The least agreement between Putzig et al. [2005] and this study occurs in class F, which is partitioned into all 10 classes in this study with at most ~ 1/5 of its pixels in any one cluster. Putzig et al.’s [2005] class F dominated the northern terrain, consisted of thermal inertia values ≥ 386 tiu with a broad range in albedo, and was interpreted as a combination of rocks, bedrock, duricrust and polar ice. The finer
265
partitioning of these pixels in this study is one of the advantages of the method of algorithmic classification and hence the high thermal inertia values have been separated into mixtures of: duricrust/pebbles/coarse sand dominated terrain (class 1); ice-cemented terrain (class 9); coarse sand/pebbles/coarse sand, dust and potential subsurface ice (classes 5 and 6).
8.4.
Summary
The validity of the 10 thermophysical classes derived in this chapter has been justified through their agreement with independent datasets such as dust cover, dune boundaries, rock abundance, extensive surface ice, measurements from surface missions, geologic units and surface age as well as visual surface morphologies.
This strongly indicates that these classes provide accurate
information on the spatial distribution of the dominant surface materials at both the global and the regional scale. The thermophysical map presented here refines previous studies on the distribution of dust, sand, duricrust, rocks and ices on the martian surface.
Both previous thermophysical maps [Mellon et al., 2000;
Palluconi and Kieffer, 1981; Putzig et al., 2005] and the map produced in this chapter, closely match in their determinations of the global spatial pattern of Mars’ fine, bright dust cover [Bandfield et al., 2000; Ruff and Christensen, 2002]. There are in general broad similarities between the map developed here and that of previous works, particularly in the equatorial regions of Mars where small unconsolidated particles dominate. However, considerable additional structure is shown here in the distribution of surface materials at mid-high latitudes, particularly in areas of medium to high thermal inertia. Furthermore, the surface material map developed here has revealed a number of trends not identified in previous publications, such as thermophysical structure in the ejecta of large martian impact craters, potential non-polar ice morphologies, and a relationship between grain size and surface age. In synthesising the results of this thesis in Chapter 9, the thermophysical map will be applied to the problem of identifying shallow, potentially hospitable environments in the martian subsurface.
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267
9. SYNTHESIS & CONCLUSIONS
‘The woods are lovely, dark and deep. But I have promises to keep, And miles to go before I sleep. And miles to go before I sleep.’ -
Robert Frost, ‘Stopping By The Woods On A Snowy Evening’
268
Synthesis & Conclusions
All known life on Earth requires liquid water and hence the search for liquid water environments outside of the Earth remains our best strategy in the search for other examples of life. This thesis has presented a holistic approach to assessing the biological potential of present-day Mars, and to identifying environments on Mars that have the highest potential of being hospitable for life. There were four main aims of this work: 1. To determine an effective way of assessing the astrobiological significance of present-day Mars. 2. To assess the potential of present-day Mars to support Earth-like life. 3. To identify candidate habitable environments on present-day Mars. 4. To suggest shallow subsurface locations which should be targeted in future missions searching for present-day life. To achieve these aims, two key complementary methodologies were applied in this thesis. The first method involved representing Earth, Mars, life, and water on pressure-temperature diagrams. Through analysing the overlap of these regions, martian environments with P-T conditions that allow liquid water, and are within the thresholds of known active life on Earth, were identified. This addressed the 1st aim and provided a methodological framework that could be applied to assess the potential of any planet to support life, assuming that any life will have similar requirements to that on Earth. The model of the potential martian biosphere that was developed provided a global assessment of habitable environments on present-day Mars and identified that extensive regions of the martian crust may be hospitable for life (addressing the 2nd aim). The second method – mapping surface materials on Mars through algorithmic classification of thermal inertia and albedo – provides constraints on the surface and shallow P-T conditions and water environments. The 10 classes of thermal surface materials derived in this work determine the shallow subsurface habitability and can be used to inform future models (also addressing the 2nd aim).
269
These methods provided complementary approaches to identifying where the pressure and temperature requirements for liquid water are met throughout current martian environments (both near surface and deeper within the crust), and what subset of these environments occur within the pressure, temperature and water activity thresholds of known active life on Earth. The latter conditions represent a set of necessary but not sufficient requirements for known terrestrial life. Section 9.1 below summarises the key results from the thesis.
Section 9.2
synthesises both approaches of this work and addresses the 3rd and 4th aim, identifying the locations on Mars with the highest potential of being hospitable to life. Section 9.3 provides a final summary of some of the applications of this research, while section 9.4 identifies potential future directions.
9.1.
Overview of results
The key results of this research are presented by topic. 9.1.1. Earth’s hydrosphere (Chapter 3) All environments on the Earth’s surface can support some form of liquid water. The coldest environments sustain liquid water with a freezing point of -89 ˚C (184 K) at the triple point of ~ 1.3×10-2 Pa. This water exists as thin films coating ice and soil grains in the Antarctic, and coating particulate aerosols up to ~ 86 km altitude within the atmosphere. The deepest and hottest possible liquid water environment, from the steepest plausible geothermal gradients, occurs at 374 ˚C and ~ 2.5×109 Pa. This corresponds to a depth of ~ 70 km and hence generally lies within the mantle. As this is the maximum possible extent of liquid water within the Earth from pressure-temperature constraints, it represents the deepest possible extent of the Earth’s biosphere, and corresponds to a hydrosphere of at most ~ 3.3 % of the volume of the Earth.
Closure of pore space and low
permeability however could conceivably limit life and liquid water environments to shallower than ~ 10-13 km in the crust, but this would vary strongly with the substrate, mineralogy and temperature conditions and is poorly constrained.
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Synthesis & Conclusions
9.1.2. The limits of life & the extent of Earth’s biosphere (Chapter 4) High temperatures provide the most significant restraint on the biosphere and its extent within the Earth.
The intersection of the 122 ˚C maximum known
temperature for life with the mean geotherm occurs at a depth of 4 km but can vary between 0-14.6 km, all well within the hydrosphere. This high temperature limit however cannot be verified as fundamental for all life as there may be a substantial amount of liquid water (< 53 % of groundwater volume or 107 km3) at higher temperatures that has not been fully searched for life.
In cold
environments, the requirement of liquid water with a water activity higher than 0.6, places a stringent restriction on hospitable environments for active life. The low volume of naturally occurring low-eutectic temperature brines may limit life on Earth to liquid water warmer than -20 ˚C. This results in 3.5-15 vol% of liquid water in permafrost soils on the Earth being inhospitable to life. All examples of life found at pressures less than ~ 10-4 Pa have been classified as dormant indicating that life at high altitude (above ~ 9 km) may be limited by severe desiccation and nutrient stress, despite the presence of liquid water droplets up to ~ 20 km altitude in the atmosphere. The apparent high pressure limit of life is likely to be a selection effect since life has been found at the maximum depths for which it has been searched. The Earth’s biosphere extends from ~ 9 km above the mean surface of the Earth, throughout the full range of ocean depths to ~ 11 km, and in at least one location ~ 5.5 km deep within continental crust and ~ 6.1 km deep within ocean and subocean sediments. This equates to the known biosphere occupying < 0.7 % of the volume of Earth. Hence after 4 billion years of evolution, active life has left > 99.3 % of Earth’s environments uninhabited, given our current exploration of subterranean life. The inhabited volume of the Earth could extend to as high as 0.9 % if active life is found in environments at 250 ˚C. The known extent of the biosphere occupies ~ 19.0 % of the volume of Earth supporting liquid water- the ~ 70 km shell of the Earth where the P-T conditions are within the liquid regime. Hence 81 % of the volume of Earth within the 70 km thick hydrosphere, is not known to harbour life. When considering the amount of liquid water, the vast majority of water on the Earth, ~ 99.2% of liquid water and ~ 47 % of groundwater, is within the temperature thresholds of life.
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9.1.3. Mars’ potential hydrosphere (Chapter 6, 7) The most plausible composition of a putative martian brine is magnesium and calcium sulphates, with additional magnesium, calcium and iron carbonates and sodium chloride, as these salts have the highest abundance in the martian soil. The total amount of salts in solution is estimated to be < 10 mol/kg or < 30 % salts by mass. Conservatively this combination of salts in solution would freeze at -10 ˚C at a triple point pressure of ~ 200 Pa. If perchlorates found at the Phoenix landing site were present in eutectic concentration within a brine, its freezing point could be as low as -69 to -75 ˚C. The coldest possible liquid water on Mars, occurring as thin liquid films in ice and permafrost, has a freezing point depression of -100 ˚C at the triple point pressure of 0.004 Pa. This broad liquid range incorporates not only thin films under the range of martian conditions, but also potentially concentrated briny inclusions such as perchlorate brines. From the mean crustal geotherm, temperatures become perennially warm enough ଷǤଽ for an average martian brine at 10.3േଵǤଷ km depth. The depth of 10.3 km, at
1.12×108 Pa, is the best estimate of the shallow, upper boundary of a potential global hydrosphere. Variations in heat flow estimates and thermal conductivities of martian materials lead to a range of 0-48.2 km for the depth of the -10 ˚C isotherm. Regional variations in insolation conditions and surface materials can allow for transient occurrences of the average martian brine at or near the surface. In particular, the warmest shallow subsurface environments occur at equatorial latitudes in low albedo, high thermal inertia materials. The deepest and hottest liquid water environment on Mars, given current geophysical models, occurs at 374 ˚C and 3.2×109 Pa, corresponding to a depth of ~ 259.4 km and hence lying within the mantle. As this is the maximum extent of liquid water, it represents the deepest, lower bound of any potential biosphere on present-day Mars. The maximum extent of the putative Martian hydrosphere corresponds to at most 21.2 % of the volume of Mars which can support liquid water-a significantly larger fraction of the planet’s volume than Earth’s. However, the existence of available pore space and the closure of rock fractures is likely to severely restrict liquid water environments below ~ 30 km depth.
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Synthesis & Conclusions
9.1.4. Potential biosphere of Mars (Chapter 6) The potential martian biosphere (PMB) was defined as all environments on Mars with pressure and temperature within the growth and metabolism thresholds of terrestrial organisms, and also within the liquid phase space of putative martian liquid water. There are many environments in the subsurface of Mars that have PT conditions that are hospitable for life. Some hot crustal environments which occur within Earth’s biosphere are absent from Mars, due to the lower geothermal heat flow and lack of present day geologic activity. Some martian environments however have P-T conditions that do not occur within the Earth and yet may be hospitable for life if liquid water is present. For example, the low mean geothermal gradients in the martian crust (mean of ~ 5 K/km compared to ~ 25 K/km on Earth) provide deep, cool environments down to ~ 10’s of kilometres that are not found on Earth. Furthermore, the lower lithostatic pressure gradient within the martian crust provides subsurface environments at pressures below the ~ 104 Pa of the lowest pressure (highest altitude) active life on Earth but within life’s temperature thresholds. The potential biosphere of Mars extends from the surface in some locations, where the requirements of: (i) temperatures above -20 ˚C; and (ii) P-T conditions within the liquid regime of water, with (iii) water activity above 0.6, are met. The depth of the potential martian biosphere is ~ 37േସଷ ଷସ km, determined by the depth at which the 122 ˚C maximum temperature for life intersects with the average modelled geotherm in the crust.
Variations in heat flow estimates and thermal
conductivities of martian materials lead to a range of 3-80 km for the lower depth limit of the biosphere, assuming the high temperature limit of 122 ˚C for life is fundamental. The most likely ~ 37 km extent of the PMB corresponds to a shell of ~ 3.2 % of the volume of the planet. Hence given the current understanding of representative conditions within the martian crust, the proportion of Mars with habitable conditions for subsurface life is five times larger than that of Earth. The ~ 37 km shell that may support Earth-like life on Mars corresponds to ~ 15.2 % of Mars with liquid water (of the volume of Mars within the ~ 260 km shell that can support liquid water). At least ~ 84.8 % of the volume of Mars that may have liquid water is inhospitable to life. This proportion of potentially hospitable liquid
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water on Mars is very similar to that of the Earth. Many subsurface environments on Mars should be hospitable for psychrophilic, halophilic oligotrophic, and anaerobic thermophillic terrestrial life, provided liquid water and sources of chemical nutrients and energy are available.
9.1.5. Distribution of martian surface materials (Chapter 8) Through algorithmic classification of martian thermal inertia and albedo 10 thermophysical units were identified and interpreted as mixtures of dust, sand, duricrust, rocks and ice on the surface. Several of these classes had not been identified in previous studies. This work found considerable additional structure in the distribution of surface materials at mid- to high latitudes compared to previous works, particularly in areas of medium to high thermal inertia. The thermophysical map reveals a pattern of concentric envelopes in the spatial occurrence of the classes in both hemispheres, reflecting underlying trends in grain size and grain cementation on the surface. Correlations were found between the interpretation of surface materials from orbital thermal inertia and albedo values, and independent observations of the surface, including: (i) the spectral dust cover index; (ii) surface morphologies and localised geologic features such as sand dunes, channel materials, impact craters, Valles Marineris and Hellas Basin; and (iii) broad terrain age. For example, the geologic units of the Valles Marineris region have a distinct thermal behaviour from their surrounding terrain and can be distinguished in the map, allowing the major canyons and the western labyrinth of valleys to be clearly defined. The borders between Noachian, Hesperian and Amazonian units are also closely matched by class boundaries in the thermophysical map. Although each class is comprised of terrain of all surface ages, there is a clear mapping from surface age to thermophysical classes. A number of large impact craters over 50 km diameter were revealed in the thermophysical map by distinct classes occurring in the crater, the crater’s ejecta blanket, and the surrounding terrain. Several classes were interpreted as sand dominated with varying dust coverage, and are strongly aligned with the spatial distribution of martian dunes > 1 km 2 with 95% of global dune area found to occur collectively within them. One class dominated the rocky region of Nili Patera and showed localised correlations to a
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Synthesis & Conclusions
number of surface features such as blocky regolith within a crater, and infilling an ancient channel. These spatial correlations strengthened the interpretation of this class as coarse cemented materials such as duricrust and/or pebbles. deposits of dust were unambiguously identified.
Thick
A number of classes are
interpreted as a mixture of dust, fine to coarse sand, and duricrust or pebbles in varying proportions throughout the horizontal and vertical spatial coverage of the pixels. One of the classes is consistent with terrain dominated by H2O ice and is correlated with: (i) the distribution of subsurface hydrogen concentrations > 50 wt%; (ii) the observations of shallow solid water ice in the northern polar terrain; (iii) the extent of the polar caps; and (iv) the interior water ice mound in Korolev crater. Where sampled directly by landers, the interpretation of each class from orbital data was consistent with the materials observed on the surface by landers. The technique of algorithmic classification was demonstrated to enable real trends in surface materials to be deciphered, and to be a highly effective tool in identifying similar pixels and hence similar surface thermal properties on Mars.
9.2.
Thesis synthesis: potentially habitable environments near the
surface of Mars. Habitable environments on Mars must meet the minimum requirements of liquid water, and water activity and temperature within the thresholds of Earth life. Thus far, water ice has been observed spectrally at both polar caps [Bibring et al., 2004; Wagstaff et al., 2008], in arctic terrain [Smith et al., 2009b], and on some other locations on the martian surface as frost/ground ice. Water ice has been inferred to be globally widespread in the top ~ 2 m of soil at latitudes above ± 50˚ [Kuzmin et al., 2009; Mitrofanov et al., 2002] and has been sampled directly at 68˚N [Mehta et al., 2011]. Liquid water has not been definitively detected on Mars, however there is strong suggestive evidence that concentrated brine was present at the Phoenix Landing site [Rennó et al., 2009] and that low volume aqueous seeps occur in other locations on the surface associated with slopes and potential subsurface volatiles [Jouannic et al., 2010; Kereszturi et al., 2010; Malin et al., 2006; Mohlmann and Kereszturi, 2010]. Water may also be present as hydrated minerals in the low latitudes below ±50˚ [Boynton et al., 2007; Milliken et al., 2007; Mustard et al.,
275
2008]. The results of this thesis provide information on the likely depth of liquid water beneath deep ice deposits, and the location of materials that provide the warmest subsurface temperatures within the 2 m depth extent of shallow ice detection- thereby allowing potential diurnal melting during the lifetime of the ice. The key factors in assessing the habitability of martian environments investigated in this thesis are summarised in Table 9-1.
9.2.1. Surface & shallow soil From the P-T model of martian environments and water chemistries, thin films are stable at any location on the surface (supported by [Mohlmann, 2008]) and concentrated brines are metastable over much of the surface depending on their specific freezing point depression (supported by [Mohlmann, 2011]).
Mars’
potential biosphere includes the surface in regions with soil temperatures above 20 ˚C, that are only likely to be encountered at the surface in low latitudes equatorwards of ~ 30˚. The duration of water will be strongly dependent on its contact with the martian atmosphere and the partial pressure of water vapour. Hydrated salty soils at ~ 10 cm depth at Gusev Crater were inferred from visual and near-infrared spectra to lose water when exposed to the atmosphere over several days [Wang et al., 2008, 2010, 2011]. Similarly, water ice buried under ~ 5 cm of soil was lost within 2 days when exposed to the atmosphere, indicating that its shallow burial depth had been sufficient to protect the ice from sublimation [Smith et al., 2009b].
These findings indicate that the shallow pore space
environment was not in equilibrium with the surface atmospheric conditions, as several centimetres of dry material significantly influenced the stability of subsurface water, increasing its lifetime.
This occurs through the thermal
insulation of the subsurface [Zent et al., 2010] and the partial protection of the subsurface from the lower partial pressure of water vapour in the atmosphere [Mellon et al., 2009a]. Below the results of this thesis will be applied to predict locations where shallow environments within the potential martian biosphere may be accessed. Examples will be given of organisms on Earth that may be able to inhabit these environments, if liquid water is present on appropriate timescales.
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Synthesis & Conclusions
Table 9-1: Factors in the selection of astrobiological targets on Mars to search for present-day life. Attribute Low elevation High humidity High salt
Recent/young flow features High water-equivalent hydrogen content Higher heat flow
Low albedo Access to subsurface Low latitude
Low thermal inertia materials (as a thin surface layer)
High thermal inertia materials
Significance High total atmospheric pressure increases stability lifetime of liquid/solid water exposed to atmosphere Increases lifetime of liquid/solid water and decreases evaporation rate; increases volume of liquid films Increases near surface humidity; increases occurrence of thin films; potentially collects water vapour onto the surface through deliquescence; lowers freezing point in aqueous solution; retards vapourisation Such as gullies and RSL’s; indicative of the presence of recent volatiles such as water Indicative of the presence of hydrated minerals or water ice/permafrost in the top 2 m; shallow water ice may be within the depth of diurnal/seasonal temperatures above melting in low albedo, high thermal inertia materials Increases the geothermal gradient within the crust and decreases the depth to melting of water ice and hospitable temperatures for life; determines temperature gradient below the ~ 20 m depth dominated by solar insolation High average surface temperatures; decreases depth to suitable temperatures for microbial life and melting of water ice Slopes such as crater walls, basins and caves allow access to potentially hospitable environments that cannot otherwise be reached without deep drilling from the surface High average surface temperatures; decreases depth to suitable temperatures for microbial life and melting of water ice; atmosphere holds higher water content at saturation Insulates deeper subsurface from surface temperature variations; small grains effectively fill pore space isolating the subsurface from the atmosphere; at low latitude the stabilisation of deeper subsurface temperature variations allows the temperatures to remain around the warm seasonal means; have the lowest temperatures overnight coinciding with water vapour saturation of the atmosphere and promoting frost and thin film deposition Provides warmest subsurface temperatures within at most the top 20 m depth (maximum depth of seasonal thermal wave); effectively transfers the minimal solar insolation through the regolith; low temperature of materials during the morning-midday promotes daily surface frost; warm temperatures during the evening high humidity promotes warm liquid film environments
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9.2.1.a.
Liquid water at the surface
Liquid films do not require the presence of ice, as they can form from water vapour condensing onto the surface. The occurrence of these films is encouraged by high humidity and high salt content [Greenspan, 1977]. Salts are present everywhere that has been sampled on the martian surface [Golombek et al., 2008b] with many of the salts expected to deliquesce under martian humidity, such as perchlorates [Marion et al., 2010; Zorzano et al., 2009]. Films of liquid can grow to a thickness of 10 monolayers (25 Å) during the warmest periods of the day [Kossacki and Markiewicz, 2010] and can persist for days without the need for atmospheric saturation [Zorzano et al., 2009], especially if dissolved salts enhance their stability [Mohlmann and Kereszturi, 2010]. At the Spirit, Opportunity and Phoenix landing sites (spanning 2˚-68˚ absolute latitude), the humidity frequently reached saturation overnight [Smith et al., 2009b; Squyres et al., 2004a, 2004b]. Thin films are therefore expected to be widespread throughout martian soil that is in contact with the atmosphere. Only in low latitude equatorial environments are thin films and brines likely to be warmer than -20 ˚C and therefore potentially hospitable for life. The warmer atmospheric temperatures at equatorial latitudes result in higher water content (max. Ppart) in the atmosphere at saturation and hence more and higher volume thin films would be expected in the soils. During the day, the warmest surface materials are those with low albedo and low thermal inertia such as fine sand. However humidity during the martian day is generally low, and hence the water content would be expected to be minimal in these low latitude materials during daytime. After sunset, low albedo and high thermal inertia materials attain their maximum temperature and are the warmest materials on the surface- coinciding with the evening peak in humidity. They are therefore predicted to have the highest content of unfrozen water as thin films and brine. Class 1, 2 and 3 terrain is dominated by dark sand with some coarse pebbles and larger rocks and occurs at low latitudes in Syrtis Major and Aram Chaos. These sites would be ideal locations to test for spectral evidence of water and hydrated salts during the martian evening, and may provide hospitable surface environments for life during warm spring and summer evenings. For example, high thermal inertia materials in southern Chryse Planitia at 25 ˚N can reach temperatures above -10 ˚C in the martian evening during summer (Appendix A,
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Synthesis & Conclusions
Table A2). The aforementioned potential surface liquid water environments are consistent with the geologic setting and temporal activity of recurring slope lineae (RSL’s), which are active in summer and observed in low albedo regions of Mars with thermal inertias consistent with sand [McEwen et al., 2011a, 2011b; Ojha et al., 2011]. This work supports an interpretation of RSL’s as thin films lubricating soil grains and reaching sufficient volume to run downslope. Furthermore, slopes in Syrtis Major and Aram Chaos are predicted to be regions of RSL formation. Flows have also been observed in low temperatures in the mid- to high latitudes [Diniega et al., 2010; Mohlmann and Kereszturi, 2010]. The conditions under which these flows occurred are inconsistent with surface melting of a representative martian brine. The perennial low temperatures (< 250 K) in these environments restricts near-surface water to either thin films, or extremely low freezing point cryobrines which have not yet been observed [Mohlmann and Thomsen, 2011]. The temperatures also restrict liquid films coating ice-free soil to being extremely thin. When ice is present in the soil, thin films will occur concurrently as a thin coating on the surface of the solid ice [Dash et al., 1995], or at the ice-soil interface [Grimm et al., 2006; Turov and Leboda, 1999]. The aforementioned cryo-flows (including dark dune spots, viscous flow features, and possibly some dune gullies), were observed in conditions below the 200 K frost point of water [Titus et al., 2003] but above the 145 K deposition of CO2 frost [Searls et al., 2010], and hence water ice was most likely present.
The low temperatures and atmospheric
saturation in the evening [Carrozzo et al., 2009; Searls et al., 2010] promote the deposition and stability of water frost on the surface. Thus ice is expected in the lowest temperature surface materials, which are those with high albedo and either high thermal inertia during the day or low thermal inertia at night. Bright fine dust would be expected to be the most ice rich as it provides the coldest surface and subsurface temperatures in the evening. Classes 7 and 8 in the thermophysical map have the highest fraction of bright dust grains < 10 µm across and are therefore predicted to show spectral evidence of water frost when temperatures are in the appropriate range. Furthermore they are predicted to be likely sites for cryo-flows, triggered by high volume thin film flow, on slopes in the regions of Utopia Planitia and southern Arcadia Planitia.
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When temperatures warm above the frost point, water ice will sublimate. The thin films of liquid associated ice can endure, with the high Ppart associated with water ice sublimation increasing their volume and their ability to flow [Grimm et al., 2006; Titus et al., 2003]. Hence although cryo-flows would predominantly occur in the ice-rich terrain at high latitudes, it is possible that they could also occur in equatorial regions in conjunction with transient water frost [Carrozzo et al., 2009; Vincendon et al., 2010a]. Surface temperatures from THEMIS nighttime infrared images indicate that high albedo, low thermal inertia materials at Gale Crater (5 ˚S) fall below 200 K in the early morning (Appendix A, Table A3). At low latitudes, classes 7 and 8 are widespread and dominate regions such as Amazonis Planitia, Tharsis and Aeolis Planum, and hence high volume thin film flow associated with transient water frost may occur in these locations. Cryo-flows however are in general unlikely to coincide with temperatures warm enough for active biology, even in the subsurface.
9.2.1.b.
Shallow subsurface liquid water
The distinction between surface and subsurface water is primarily in its duration. Subsurface brines can occur from the melting of transient ice.
Water ice is
abundant in the soil at latitudes above 50˚ in both hemispheres, however only at lower latitudes will the subsurface soil temperatures reach above the -20 ˚C for life or the -10 ˚C melting point of a representative martian brine. There is growing spectral evidence for fall-to-early spring deposition of frost at latitudes polewards of 13˚ in the south and > 32˚ in the north as layers 20-50 µm thick [Carrozzo et al., 2009; Vincendon et al., 2010a]. Frost was also observed at 13 ˚S and 2 ˚N by Spirit and Opportunity, respectively [Schorghofer and Edgett, 2006]. Modelling suggests that bright, high conductivity (clean) water ice over dry soil will generate transient liquid water at the ice-soil interface in warm temperatures, at any location on the martian surface [Hecht, 2002]. For example, shallow (~ 10 cm) melting beneath ice may be responsible for the pure liquid phase inferred to have been present at the Phoenix Landing site [Rennó et al., 2009; Smith et al., 2009b]. However, any absorbed pure or briny liquid water exposed to the atmosphere will be metastable and lost in a matter of days [Haberle et al., 2001; Squyres et al., 2004a]. Shallow ice effectively plugs pore space and isolates the lower subsurface from the martian atmosphere, as well as conducting solar heat throughout its bulk to warm the
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Synthesis & Conclusions
subsurface environment.
Hence, equatorial seasonal ice environments can
promote the temporary stability of liquid water beneath the ice, until the ice plug is lost through sublimation. Given the above discussion, it is plausible therefore that subsurface water ice could melt within the top ~ 3 m- the seasonal skin depth in icy soil- producing transient liquid water in environments with the warmest subsurface conditions. This is consistent with recent early spring flows in mid-latitude gullies [Dundas et al., 2010, 2011a; Malin et al., 2006; Pelletier et al., 2008; Reiss et al., 2010] and with modelling of gully alcove environments [Heldmann et al., 2005; Kossacki and Markiewicz, 2004]. For example, recent gully activity occurred in the Centauri Montes region (38 ˚S) within class 4 or 5 terrain [Dundas et al., 2010], and Russell crater (54 ˚S) within class 2 or 3 terrain [Reiss et al., 2010]. These new flows generally align well with the warmest subsurface environments, which occur beneath the low albedo and high thermal inertia materials that dominate class 1-3 (and class 4 to a lesser extent) terrain. These thermophysical units are also coincident with moderate hydrogen content, > 5 wt%, within the top 2 m in the regions of Mawrth Vallis, Oxia Palus and eastern Meridiani Planum [Boynton et al., 2002a; Mitrofanov et al., 2002].
Hence these regions may be ideal sites to
investigate the early-spring subsurface melting and runoff of late winter water ice deposited through the shallow regolith in high thermal inertia materials. Furthermore, low- and mid-latitude gullies may fall within the potential martian biosphere.
9.2.2. Organisms with relevant adaptations for Mars 9.2.2.a.
Surface environments
The water activities of the aforementioned liquid environments in section 9.2.1 are likely to be low due to subzero temperatures and high concentrations of solutes. Hence they may be predominantly beyond the desiccation thresholds of terrestrial biology (aw = 0.6). For example, a water activity of 0.5 is likely for liquid brines consistent with viscous dune flows at 180 K [Mohlmann and Kereszturi, 2010]. It is plausible however that some combinations of chlorides, sulphates and carbonates
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could maintain a subzero brine activity above the 0.6 threshold for life, providing potentially viable habitats for extremophiles. In Earth permafrost, bacterial cells are covered by thin films of water that allow for nutrients and waste exchange via diffusion in the films [Rivkina et al., 2000]. Studies of methanogens from Siberian permafrost in simulated pressure, temperature, atmospheric composition and water vapour conditions of the martian surface in the mid-latitudes, found a 90% survival rate [Morozova et al., 2007]. The methanogens utilised liquid water in the soil from adsorbed water vapour with a variable activity between 0.1-0.9. This study and others show that the adaptations of microbes that survive in permafrost and are exposed to desiccation and low temperature stress, are well suited to the large diurnal variations in temperatures and water activity, and the transient liquid water availability consistent with the shallow martian subsurface [Amato et al., 2010]. Cyanobacteria in the Atacama Desert harvest water vapour that deliquesces onto hygroscopic (able to attract moisture from the atmosphere) surface minerals during high humidity [Davila et al., 2008; Wierzchos et al., 2011]. Other organisms on Earth, such as lichens and lithoautotrophic bacteria, can withstand the conditions in the shallow subsurface, if they are able to employ the same strategy of utilizing thin films and briny droplets for water [Jepsen et al., 2007; de Vera et al., 2010]. However, the high UV flux on the martian surface would sterilise much life within the top ~ 1 cm [Peeters et al., 2010; Schuerger et al., 2006], with the exception of some radiation resistant bacterial spores [Dartnell, 2011; Dartnell et al., 2011; Kerney and Schuerger, 2011; Nicholson et al., 2005; Schuerger and Nicholson, 2005]. Beneath this depth however, many terrestrial organisms could survive the UV and temperature stress, if sufficient liquid, energy and nutrients are available [Diaz and Schulze-Makuch, 2006; Gómez et al., 2010].
9.2.2.b.
Subsurface environments
Unlike surface life, life in subsurface envrionments wil be protected from sterilising UV flux, and is shielded to some degree from the large seasonal temperature variations of the surface.
It is also excluded from the other
wavelengths of light, and hence cannot derive energy through photosynthesis. From the martian phase model, the majority of Mars (represented by the mean
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Synthesis & Conclusions
geotherm) is below the minimum temperature for life of -20 ˚C at depths shallower than ~ 9 km. Psychrophiles in sub-zero liquid water will also face challenges of low activity intracellular fluid and slow diffusion gradients [Piette et al., 2011]. At temperatures above -20 ˚C, terrestrial psychrophiles and halobacteria that grow in low temperature, high salinity, and desiccated environments (such as permafrost and ice) may be well suited to the martian subsurface, assuming liquid water is present [Horneck, 2000, 2008; Landis, 2001; Mancinelli et al., 2004; Pavlov et al., 2009]. One example of an organism well adapted to the martian subsurface is bacteria found 2.75 km deep within an Arctic ice core at -20 ˚C [Junge et al., 2004]. In deeper environments, above 0 ˚C, within putative liquid brine aquifers, the P-T conditions on Mars would be analogous to those faced by deep life on Earth [Pedersen, 2000, 2010]. Hence these environments may be habitable for anaerobic bacteria, such as organisms found within fluid fractures in a gold mine 2.8 km beneath the surface [Chivian et al., 2008], if energy sources are available on Mars [Boston et al., 1992; Sherwood Lollar et al., 2007].
9.3.
Conclusions
Throughout this thesis a broad framework for assessing the biological potential of different solar system environments through pressure and temperature mapping was developed.
A more detailed and accurate global mapping of dust, sand,
duricrust, rocks and ice on the martian surface was produced. Furthermore, important information on martian surface materials was derived, which can be utilised in two key ways: firstly, to assess the near-surface habitability, as demonstrated by the subsurface temperature modelling of materials surrounding Gale Crater; and secondly, to inform the assessment of landing site suitability through the identification of surface hazards (such as thick dust and other low traction materials, and high rock abundance). This research provided a rigorous assessment of the current pressure and temperature limits of life on Earth focusing on whether the observed limits are real- and hence should be applied when determining the biological potential of other environments- or are a selection effect based on our limited exploration of the Earth.
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Ultimately, it was shown that Mars has an extensive range of environments potentially hospitable to life in the crust below 10 km depth, where liquid brines can be perennially stable.
Thin films can occur perennially throughout
environments that are in diffusive contact with the martian atmosphere or with water ice. Brines can exist transiently at the surface and throughout the soil, however at high latitudes the eutectic temperatures must be extremely low (requiring a high concentration of salts) to maintain the liquid phase. Although brines will be metastable when exposed to the atmosphere, the factors which promote fluid formation (high humidity and salts) also lengthen its lifetime and enable liquid water to occur transiently on diurnal timescales across the surface. At low latitudes, and in low albedo and high thermal inertia materials that provide the warmest shallow subsurface environments, liquid water can coexist with temperatures within the thresholds of Earth life in the top ~10 m of regolith. These environments are within reach of the next decade’s robotic explorers of the martian surface. Syrtis Major, Oxia Palus, Mawrth Vallis and eastern Meridiani Planum have a high potential for hospitable liquid water environments at the surface and in the shallow subsurface, < 10 m depth.
At high latitudes, the
prevailing low temperatures drive the minimum depth for life deep within the crust. This work has provided significant results to aid the selection of sites with astrobiological significance within the shallow martian soil and hence in reach of the next generation of explorers. The identification of near-surface environments on Mars that have the potential for liquid water is relevant to NASA’s Mars Science Laboratory (MSL) mission scheduled to arrive at Gale Crater in August 2012 [Grant et al., 2011]. MSL will assess the habitability of its landing site through: (1) determining the nature and inventory of organic carbon compounds; (2) determining the molecular configuration, oxidation state and isotopic composition of the chemical building blocks of life; and (3) determining the present state, distribution and cycling of water and its persistence at the landing site [McCleese, 2003].
Through a neutron spectrometer (the Dynamic Albedo of Neutrons
package) MSL will have the capability to measure the concentration of adsorbed water (H2O) and hydrated minerals (OH) in the top ~ 1 m of martian soil [Busch and Aharonson, 2008; Litvak et al., 2008]. This mission will therefore have the
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Synthesis & Conclusions
capability to test a number of the predictions made in 9.2.1 above. Specifically, whether dust and sand within MSL’s predicted landing ellipse is cemented with water ice during high humidity. The potential martian biosphere model also identifies planetary protection regions [Chyba and Hand, 2005]. If environments within the potential martian biosphere were contaminated by life on Earth, particularly those organisms discussed in 9.2.2.a, the conditions could allow their survival and potential growth if liquid water with an activity above 0.6 is present. Environments identified as lying within the potential martian biosphere, such as the locations of recent flows in gullies and RSL’s discussed above, should be added to the list of currently designated special regions to prevent forward contamination [Beaty et al., 2006; Chyba et al., 2006].
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9.4.
Future directions
Some of the areas into which the work presented in this thesis can be expanded are discussed below.
9.4.1. Habitability of Mars through time Throughout this work, several mentions have been made of the changing stability of liquid water on the martian surface through time, driven by the oscillations of Mars’ obliquity. Mars’ obliquity undergoes large variation between extremes of 10˚-60˚ in cycles of ~ 106 years [Laskar, 2004]. These variations have had a significant effect on the martian climate, with the higher obliquity during the Noachian (> 3.5 Ga ago) resulting in higher insolation at the martian poles and a thicker atmosphere due to the higher vapour pressure of CO2, which in turn led to a greenhouse effect and moderately higher global surface temperatures [Johnson et al., 2008]. The surface of Mars indicates a very different liquid water environment in the past, with evidence of significant fluvial erosion through valleys, outflow channels, a northern ocean [Mouginot et al., 2012] and numerous lakes [Cabrol and Grin, 2002] on surfaces older than ~ 500 Ma [Jakosky and Phillips, 2001; Page et al., 2009].
Although the cause of significant liquid water discharge is not well
understood [Moore et al., 2003], the timescales of liquid water activity can be partially constrained and used to develop a model of the extent of habitable environments on early Mars.
This provides motivation for future work on
modelling the evolution of the martian pressure-temperature phase space and the potential martian biosphere through time. Other factors that have varied with Mars’ obliquity, such as the surface UV radiation environment, could also be incorporated to constrain the habitability of near-surface environments for past life [Ehresmann et al., 2011]. Temporal variations in the pressure and temperature conditions on Earth could also be included into the Earth phase space model, to investigate the evolution of the Earth’s habitability through time and constrain the duration of habitable environments.
286
Synthesis & Conclusions
9.4.2. Further characterisation of thermophysical units Studies into the mineralogy of the martian surface have revealed some correlation between certain minerals and the albedo and thermal inertia of the surface materials. Dark soils on Mars are typically sulphate poor (< 5%) [Gendrin et al., 2005; Langevin et al., 2005a; Yen et al., 2005] with sulphates predominantly occurring in bright fine-grained materials [Bishop et al., 2005; Farrand et al., 2009; Foley, 2003; Wang et al., 2010]. Clays (phyllosilicates) are typically found in dark deposits, as are pyroxenes [Ehlmann et al., 2008; Poulet et al., 2005]. Hence the occurrence of some minerals may correlate with the thermophysical map, allowing the classes to be further characterised in terms of the dominant mineralogy and providing information on the variation of mineralogy with region and geologic history. The correlation between sulphates and surface materials is particularly significant as sulphates enhance the preservation of biological material, and hence could preserve a record of life if any existed on Mars [Aubrey et al., 2006; Skelley et al., 2007]. This discussion presents the motivation for future work in combining the thermophysical map with visual imagery and mineral spectra to further define the key surficial materials on Mars and how they were emplaced. Two applications are given below.
9.4.2.a.
Tracing the origin of fines
Fines on Mars can be derived from both: (i) chemical processes, such as oxidation and interaction with present-day thin liquid films and with extensive liquid water (and hydrothermal environments) in the martian past [Baird and Clark, 1981; Burns, 1993; Gibson et al., 1983; Golombek and Bridges, 2000; Yen et al., 2005]; and (ii) mechanical processes, such as aeolian weathering (abrasion), impact cratering, and freeze-thaw cycles degrading rock [Bibring et al., 2006; Gooding et al., 1992]. Palagonite
indicates
a
chemical
weathering
origin
of
fines
(through
palagonitisation of basalt [Bell et al., 2000; McSween Jr. and Keil, 2000]). Correlating mineral maps of palagonite, with grain size from the thermophysical classes derived in Chapter 8, will allow for a more conclusive interpretation of whether chemical or mechanical weathering has been dominant in different regions and for different materials on the martian surface [Poulet et al., 2007].
287
9.4.2.b.
Understanding the effect of past water activity on grain size
Given the efficacy of the thermophysical map in discriminating different geologic units, particularly those dominated by larger grain size- such as the examples given in Chapter 8 of coarse channel and blocky crater fill, and coarse and icy interiors of Lomonosev and Korolev craters- the thermophysical classes may prove useful in identifying the sorting, rounding, and degradation of surface materials in relation to both ancient episodes of fluvial activity and water freezing [Burr et al., 2002; Hall and Andre, 2003].
288
Synthesis & Conclusions
289
‘...I want you, and that’s so terrifying And I want you, to help put out the fire ‘Cause I am an island, and you are the ocean, and all of my sadness taken by the sea. And I want so much to believe that I won’t disappear in the water. That I won’t always be swimming against the tide.’ - Darren Hayes, ‘Taken By The Sea’
290
291
APPENDIX A: IMAGE METADATA This appendix contains the metadata for all HiRISE and THEMIS images that were used for illustration or analysis within the thesis. The metadata fields are a selected subset of the available data, chosen for scientific relevance.
Solar
longitude indicates the martian season in which the image was taken; local time indicates the time of day (24hr system). Together these parameters indicate the insolation environment of the surface.
Where available the range of surface
temperatures and thermal inertias within the image is also given. In the tables below, solar longitude is coloured to indicate season in the respective hemisphere according to the following: Summer Autumn Winter Spring
Table A1: Assorted HiRISE red wavelength images Image ID ESP_011396_1115 ESP_011883_1325 ESP_012024_1440 ESP_013017_1325 ESP_014093_1410 ESP_015918_1325 ESP_016907_1330
Solar longitude 183.24 205.50 212.17 260.77 312.06 25.55 60.10
Local time
Max. latitude
16.30 15.96 15.94 15.22 14.27 14.73 15.25
-68.35 -46.85 -35.63 -47.13 -38.68 -46.82 -46.54
Max. longitude 1.47 19.69 129.50 19.56 159.58 18.89 20.24
292
Appendix A: Image Metadata
Table A2: Assorted THEMIS images. Image ID’s commencing with ‘I’ indicates infrared wavelength, ‘V’ indicates visual wavelength. Image ID
I00850009 I00880002 I00905006 I00931002 I01524004 I01886009 I02660002 I03359002 I03658006 I04407005 I04482005 I04507005 I04526006 I04919006 I06461011 I08215018 I09042008 I11073008 V10449014 V14742009
Solar Min. longitude local time 331.17 14.99 332.50 3.25 333.62 3.25 334.79 3.03 0.17 15.36 14.71 15.55 44.01 16.06 69.32 16.43 80.07 16.85 107.45 17.11 110.26 17.14 111.20 17.14 111.92 17.31 127.03 17.25 194.14 5.35 283.96 16.50 324.16 4.00 46.51 16.89 23.17 16.61 191.11 17.72
Central latitude 8.76 -46.39 -45.42 4.19 20.06 17.57 22.22 22.23 64.46 64.26 64.72 64.27 72.52 65.47 0.91 20.02 21.06 22.62 23.22 24.13
Central Max. longitude surface temp. 323.79 285.79 19.03 206.30 17.87 208.36 340.67 218.72 326.72 275.95 326.16 278.16 325.26 269.55 326.53 262.67 352.39 249.69 353.19 252.93 350.63 264.59 349.55 255.39 165.72 245.55 348.87 250.81 341.63 218.63 326.07 248.70 62.49 205.75 325.79 251.07 326.14 326.44
Min. surface temp. 216.70 178.09 185.74 187.57 223.77 224.08 219.90 217.46 183.69 195.37 198.49 182.21 153.93 181.85 174.06 189.55 178.88 217.47
Table A3: THEMIS infrared images of Gale crater and surrounding region Image ID
Solar Min. longitude local time
Max. thermal inertia
Min. Max. surface Min. surface thermal temperature temperature inertia
I18262008 I34958006 I09938003 I18337010 I18362023 I01687003 I35120006 I35145006 I35170005 I35195003 I35220022 I18549006 I35245009 I18574010 I18624012
2.15 2.55 2.97 5.20 6.21 6.78 9.09 10.08 11.08 12.07 13.06 13.67 14.05 14.65 16.61
1015.8 552 659.9 703.9 580.6 921.4 1128.8 753.9 927.9 731 495.1 932.9 618.9 875.1 946.9
83.1 80.9 83.1 88.7 118.9 104.5 97 105.7 112.7 70.1 133.9 90.3 86.9 88.3 21.7
4.43 3.39 4.34 4.43 4.45 3.47s 3.43 3.43 3.43 3.45 3.44 4.57 3.47 4.59 4.55
211.88 203.04 205.09 204.87 204.22 210.64 211.57 202.92 210.51 205.95 199.01 207.75 202.42 208.19 208.52
163.84 166.96 165.78 167.25 168.00 170.00 166.98 167.45 170.26 163.40 173.49 164.47 165.80 161.32 152.09
293
I27010015 I10375006 I27085019 I27110019 I35457009 I35482011 I18811004 I35557005 I27322003 I35694002 I27372009 I35719006 I35744006 I35794006 I35819006 I27584023 I35931007 I35981004 I27659007 I36006006 I36031005 I36056005 I27734006 I36081005 I27759006 I36106007 I19460021 I11161022 I11211003 I36243009 I11286022 I28021005 I28046005 I36418012 I36443025 I03160022 I36555023 I36605004 I36630004 I03260022 I36680005 I36705003 I36730007 I28520022 I37154002 I37254007 I37279007 I37329002 I37354002
18.26 20.30 21.17 22.13 22.30 23.26 23.85 26.13 30.21 31.32 32.10 32.26 33.20 35.07 36.00 40.00 40.16 42.01 42.77 42.93 43.85 44.77 45.53 45.69 46.45 46.60 48.08 49.71 51.53 51.60 54.25 55.98 56.88 57.94 58.85 62.15 62.88 64.68 65.58 65.75 67.38 68.28 69.18 73.95 84.44 88.06 88.97 90.78 91.69
4.78 4.58 4.82 4.83 3.63 3.57 4.67 3.55 5.00 3.60 4.98 3.62 3.62 3.61 3.62 5.09 3.65 3.67 5.14 3.69 3.70 3.70 5.16 3.72 5.17 3.71 4.97 4.91 5.03 3.83 4.96 5.29 5.30 3.84 3.83 4.33 3.89 3.92 3.93 4.38 3.97 3.94 3.98 5.45 4.07 4.07 4.07 4.16 4.16
550 1166.9 709.4 867.1 1367.9 727.8 464.3 1069.3 2000 737.4 670.8 838.6 860.9 2000 2000 510.7 2000 2000 779.1 2000 2000 689.8 625.5 2000 1330.1 900.5 346.7 388.2 2000 2000 731.7 2000 815.7 2000 1156 564.5 460 2000 2000 539.4 2000 2000 2000 667.2 2000 1241.5 695.5 2000 2000
31.8 112.9 50.3 48.6 120.9 61.9 20.4 76.7 66.8 103.5 63.8 91.8 65.5 102 42.4 113 45.5 72.7 75 88.5 72.5 59 46.8 109.7 53.3 64.8 67.1 118.2 118.6 101.3 100.5 54.8 58.9 48.4 109.5 132.4 88.8 96.1 90.3 92.5 138.6 77.9 86.1 110.6 43.4 76.9 41.8 136.3 104.2
199.48 211.50 203.26 207.38 203.26 202.94 195.23 203.88 199.46 200.83 200.37 199.30 200.43 202.77 199.24 191.03 197.75 198.34 199.89 196.01 200.62 200.56 195.61 197.70 206.83 204.01 183.86 186.93 205.25 199.90 197.90 196.58 200.96 196.93 202.45 194.06 185.52 197.99 200.70 192.18 201.61 203.08 195.85 193.82 191.82 200.84 195.39 198.05 199.51
153.90 164.80 158.08 158.41 163.77 160.92 150.45 161.67 150.74 166.69 157.40 164.76 156.20 158.45 147.83 162.90 149.45 150.81 153.64 154.89 152.84 154.00 151.99 158.44 153.27 156.17 155.54 165.56 154.31 148.55 162.78 152.69 148.89 151.79 163.64 169.01 157.16 153.21 156.96 160.80 159.81 155.99 149.17 163.52 150.12 158.61 151.89 150.36 152.49
294
I37441006 I37466011 I37591009 I29281022 I29331002 I04346022 I29381002 I29431002 I37778022 I37853008 I29618002 I29668002 I29693002 I13108022 I29880008 I29930002 I29955002 I04945022 I38377002 I05020011 I05195002 I22006023 I38764002 I38789002 I38814002 I30504022 I05507012 I39026003 I05694011 I14156022 I05844009 I14206004 I05869008 I05919007 I14281006 I31115011 I31140009 I39600022 I06306015 I31402004 I06393017 I06443015 I14805013 I06468013 I31514009 I14830009 I31564008 I06568014 I06593010
Appendix A: Image Metadata
94.87 95.78 100.38 101.59 103.44 105.16 105.30 107.16 107.32 110.13 114.20 116.11 117.06 120.97 124.28 126.23 127.21 128.04 130.34 131.01 138.03 143.07 146.05 147.10 148.14 149.52 150.95 157.16 158.97 164.12 165.55 166.33 166.66 168.90 169.69 176.39 177.54 182.97 186.73 189.83 190.87 193.28 194.12 194.48 195.22 195.33 197.65 199.35 200.57
4.16 4.14 4.18 5.64 5.66 4.78 5.69 5.68 4.17 4.26 5.72 5.73 5.77 5.41 5.77 5.79 5.80 4.98 4.30 5.02 5.08 5.66 4.44 4.45 4.39 5.72 5.24 4.46 5.28 5.63 5.32 5.74 5.33 5.33 5.70 5.52 5.51 4.51 5.44 5.42 5.44 5.45 5.74 5.45 5.34 5.73 5.30 5.47 5.45
2000 2000 2000 299.6 607.6 343.7 2000 2000 313.8 2000 2000 1104.7 2000 531.5 853.7 2000 618.4 260.4 533.6 1251.8 1357.7 598.7 2000 2000 673.4 431.9 899.9 681.8 665.6 224.2 1035.6 1277.3 926.7 911.6 636.7 758.8 686.9 399 716.2 630.5 520.3 725.8 549.2 547.2 492.3 542 471.5 707.4 654.5
94 93.1 129.1 56 52.3 91.5 97 156.2 50.5 117.6 82.1 210.4 240.6 138.1 86.4 89.5 70.8 62.6 87.3 74.1 85.2 151.6 156.5 176.8 159.3 98.1 81.9 78.6 69.2 79.4 176.8 75 58.2 119.4 66.8 136.9 75.1 133.1 64 89.8 43 64.8 104.6 37 54.4 81.9 97 101.7 67.3
192.77 198.93 196.69 177.31 189.80 181.91 198.09 198.97 180.77 200.22 192.81 202.65 202.91 191.57 197.78 197.80 196.60 179.17 195.62 203.53 210.70 194.49 202.66 206.42 203.25 191.77 203.52 202.53 201.01 180.66 211.74 205.52 211.79 212.79 204.33 208.67 207.54 198.71 208.63 206.15 203.32 210.54 213.37 205.39 201.95 204.97 202.61 211.38 209.58
149.56 153.42 150.08 149.32 150.50 157.29 153.96 164.64 150.73 153.75 155.26 171.83 167.98 165.61 158.48 154.60 155.47 153.66 160.66 156.50 161.73 166.03 165.66 168.30 174.88 162.01 155.89 157.45 160.82 160.73 177.24 154.95 156.90 171.41 157.85 174.86 164.80 175.69 163.51 168.60 157.94 163.68 168.11 158.83 160.78 161.79 170.74 168.45 165.11
295
I06643009 I40049009 I40074006 I31801015 I06780009 I06805007 I06855006 I31901010 I06880008 I31926006 I15242023 I06955008 I31988009 I07005012 I32163006 I07142011 I07167013 I07192006 I32225013 I07217007 I32250010 I07242018 I07267013 I32300002 I07292010 I32325004 I07317010 I07342010 I07367010 I07392005 I07417014 I07504013 I07529010 I07554003 I07579011 I07604009 I07629023 I07654010 I07679011 I07704010 I32749015 I07729010 I07754014 I07779011 I41160023 I32824005 I32849003 I07891012 I07916015
203.03 204.62 205.85 209.36 209.84 211.10 213.61 214.39 214.88 215.65 215.76 218.68 218.80 221.23 227.77 228.27 229.56 230.85 230.97 232.14 232.27 233.44 234.73 234.86 236.03 236.16 237.33 238.63 239.94 241.24 242.54 247.09 248.39 249.70 251.01 252.31 253.62 254.93 256.23 257.54 258.29 258.84 260.14 261.45 261.81 262.20 263.51 267.28 268.57
5.52 4.51 4.53 5.15 5.49 5.50 5.45 5.06 5.45 5.03 5.59 5.43 4.98 5.44 4.81 5.35 5.35 5.32 4.74 5.31 4.72 5.33 5.32 4.64 5.26 4.62 5.28 5.23 5.27 5.22 5.24 5.20 5.21 5.12 5.13 5.09 5.04 5.05 5.04 5.04 4.08 4.99 5.04 5.02 3.82 3.94 3.91 4.88 4.81
744.5 541.5 807.4 528.6 688.5 623.1 577.8 616.2 666.1 857.5 573.6 820.4 457.3 822.9 813.7 459.8 495.1 704 570.5 587.2 627 877.6 776.5 504.6 891.7 614.2 918.3 560 635.5 646.8 682.6 638.9 598.3 648.1 655.5 708.6 604.4 671.9 866.9 732 670.8 751 741.3 642.3 630.1 611.1 639 757.2 2000
114.6 80.6 62.9 94.2 65.9 104.9 63.4 99.2 110.8 114.1 77.4 181.8 101.9 129.1 115.4 131.1 79.4 90.5 152.8 104.6 145.7 138 187.9 153.1 126.2 123.2 111.3 125.9 91.1 106.1 78.1 130 113.6 86.6 125.5 117.9 129.2 89.4 104.9 91.2 140.5 100.5 103.3 107.3 140.9 105.6 138.5 102.7 113
211.59 208.11 215.63 207.23 210.10 216.42 207.89 211.10 211.39 218.92 208.11 215.67 207.13 218.05 218.75 207.64 205.82 213.65 211.35 207.49 212.03 217.24 219.42 208.49 214.38 213.59 216.69 208.88 213.71 213.05 218.27 223.08 223.66 208.56 216.71 213.98 212.45 214.94 217.67 220.78 218.96 216.09 220.43 218.20 214.73 214.99 215.07 218.82 208.79
173.28 168.43 164.17 170.23 161.65 169.46 166.61 173.30 175.30 175.93 169.00 179.98 173.00 177.77 177.39 177.33 170.58 167.05 180.62 171.40 182.87 178.63 186.93 181.46 175.76 182.15 176.68 177.98 170.36 176.48 170.03 178.06 175.76 171.75 177.01 177.13 179.32 175.85 175.44 173.66 181.48 174.75 175.34 174.90 181.51 179.45 180.97 173.45 164.10
296
I07941013 I07966017 I07991013 I08091015 I08116012 I08141016 I24926022 I08253024 I33298005 I08278013 I16702023 I25051022 I33473005 I08453006 I33585006 I33610010 I33685006 I33710006 I17114009 I08790008 I33922006 I17251013 I25650026 I34022006 I34047006 I17376009 I25862005 I34209006 I17563023 I17588013 I34284008 I17613014 I17663019 I00988002 I17825025 I34521006 I34571007 I09551010 I17925025 I34621022 I17950012 I01325006 I01350002 I26436014 I34783006 I34808006 I18162009 I34858008 I34883005
Appendix A: Image Metadata
269.87 271.16 272.45 277.61 278.89 280.16 285.10 285.87 286.60 287.13 291.14 291.40 295.39 295.91 300.93 302.15 305.81 307.02 311.36 312.33 317.15 317.86 320.44 321.83 322.98 323.68 330.18 330.40 332.21 333.33 333.77 334.45 336.68 337.32 343.78 344.20 346.36 346.81 348.08 348.50 349.15 351.87 352.92 355.14 355.33 356.37 358.04 358.45 359.48
4.79 4.81 4.79 4.64 4.67 4.67 4.73 4.50 3.54 4.55 4.49 4.64 3.37 4.46 3.33 3.33 3.26 3.25 4.28 4.20 3.19 4.23 4.39 3.23 3.22 4.22 4.38 3.19 4.20 4.24 3.19 4.23 4.23 3.15 4.29 3.29 3.24 4.21 4.28 3.25 4.29 3.28 3.28 4.53 3.36 3.37 4.39 3.35 3.35
687.3 667.5 746 713.5 683.5 702.3 434.5 301.2 898.2 742.3 618.1 459.5 681.9 686.7 996.3 853.5 836.4 1010.3 727.8 550.1 832.9 739.1 612.2 978.4 821.6 585.8 711.2 856.9 498.5 682.4 1079 685.2 967.6 809.2 751.7 892.2 943.7 508.2 608.8 612.2 840.1 916.8 776.6 982.6 883.7 1270.1 848.8 1051.9 823.2
115.9 122.7 87 89.6 63.4 109.9 282.3 98.5 135.9 98.7 121.1 243.7 129.9 109.6 92.2 143.8 159.2 113.5 107.6 86.3 117 96.2 188.8 114.1 113.9 108.9 108.8 209 134.9 97.5 127.2 96.6 111 136 108.9 164.7 150.3 103.9 122 165.4 103 120.6 94.8 81.9 126.3 155.2 64.1 111.9 116.3
212.43 214.42 214.19 211.58 217.37 217.81 202.34 195.00 220.74 217.92 211.81 203.82 215.25 218.44 215.75 215.54 215.31 216.96 212.07 206.40 216.59 209.35 208.09 218.86 216.04 207.92 208.01 213.30 202.47 208.18 213.25 208.25 211.09 214.42 208.17 211.31 215.30 202.68 204.35 206.16 212.03 213.58 211.23 211.17 209.80 215.72 207.32 214.14 211.02
171.54 176.86 174.63 170.62 167.36 177.34 193.26 172.51 178.95 171.50 177.04 190.32 181.18 174.61 170.34 179.89 182.39 176.05 173.05 170.43 176.54 173.33 183.84 175.43 176.78 170.39 173.85 185.32 175.14 167.49 176.45 167.16 169.19 179.31 169.14 179.10 178.66 168.78 171.60 179.48 170.77 171.71 170.21 166.28 168.86 174.07 159.18 171.35 173.77
297
APPENDIX B: PROCESSING PROCEDURE FOR HIRISE & THEMIS IMAGES This appendix contains the processing methodology that was undertaken for all HiRISE and THEMIS images. Pre-processing of the images by the instrument teams, in the form of radiance calibration, stitching and geometrically projection, was undertaken before the images were made available.
HiRISE images The images from the High Resolution Imaging Science Experiment (HiRISE) utilised in this study were sourced from http://hirise.lpl.arizona.edu/. Only single band images, called ‘red’ images (although displayed in greyscale) were used. These images give the surface reflectance in the visual wavelength range of 550850nm (towards the red end of the spectrum)
[Eliason et al., 2009].
The
maximum image size is 20,000×126,000 pixels, with a resolution of 0.25-0.5 m/pixel. Further details on calibration that occurs during .JP2 image production can be found at [Becker et al., 2007; Delamere et al., 2010].
Processing i.
Download .JP2 (jpeg2000 image file) and .LBL (metadata file) to the same directory
ii.
Within ISIS3, run the perl procedure ‘pds2world.pl’ on the .LBL files with the following command: for i in *.LBL; do perl pds2world.pl –J –prj $i; done The –J command produces .jpw world files and the –prj command produces projection files. Together they contain the spatial information associated with each image that ArcGIS requires to geo-reference and project the images.
iii.
Ensure that the naming of .jpw and .prj files matches the images, and they are placed in the same directory
298
iv.
Appendix B: Processing Procedure for HiRISE & THEMIS Images
Within ArcGIS, run the USGS Image Toolbox v1.5 tool ‘Batch set image projection from PRJ’. This directs ArcMap to read the spatial information for each image from the .prj file.
v.
Import the .jp2 files into ArcMap with the setting ‘Use world file to define the coordinates of the raster’ unchecked.
THEMIS images The images from the Thermal Emission Imaging System (THEMIS) utilised in this study were identified through Java Mission-planning and Analysis for Remote Sensing system (JMARS: http://jmars.asu.edu/). Although the images contained multiple wavelengths, only band 9 was used. Band 9 images give the surface radiance in the infrared wavelength range of ~12-13 µm (centred at 12.57µm) [Christensen et al., 2004c]. Each image is typically 1200×3700 pixels, with a resolution of 100m/pixel [Edwards et al., 2011]. Further details on calibration that occurs during .TIFF image production can be found at [McTiernan et al., 2009; Murray, 2010] and in the automated THEMIS processing website (THMPROC) http://themis.asu.edu/thmproc.
Processing i.
Identify image ID’s from JMARS with the following constraints: - Summing: 1 to 1 - Image rating: 6 to 7 - Band 9 available
ii.
Request processing of images through THMPROC with input parameters: - Unrectify - SIMP, 0, OCENTRIC, -180/180, 0.1 - 32 bit band 9 radiance cube (keep backplanes unchecked) - 8 bit band 9 TIFF, stretch
iii.
Download .TIFF (image file) and .QUB (metadata file) to the same directory when processing is completed
iv.
Within ISIS3, run the perl procedure ‘isis2world.pl’ on the .QUB files with the following command: for i in *.QUB; do perl isis3world.pl –t –prj $i; done
299
The –t command produces .tfw world files and the –prj command produces projection files. Together they contain the spatial information information associated with each image that ArcGIS requires to geo-reference and project the images. v.
Ensure that the naming of .tfw and .prj files matches the images, and they are placed in the same directory
vi.
Within ArcGIS, run the USGS Image Toolbox v1.5 tool ‘Batch set image projection from PRJ’. This directs ArcMap to read the spatial information for each image from the .prj file.
vii.
Import the .tiff files into ArcMap with the setting ‘Use world file to define the coordinates of the raster’ unchecked.
viii.
Separate daytime and nighttime images. Daytime images usually slant NESW and appear as photographs, nighttime slant NW-SE and appear fuzzy.
300
Appendix B: Processing Procedure for HiRISE & THEMIS Images
301
APPENDIX C: METHODOLOGICAL CONSIDERATIONS FOR CHAPTER 8 Optimal number of thermophysical units It was discussed in Chapter 8 that the clustering algorithms used provided an extremely useful partitioning of global martian thermal inertia and albedo into 10 classes, as evidenced by the strong correlation between class properties and class boundaries with independent datasets on martian surface materials (such as spectral properties, geologic units, surface age and surface morphology from visual and infrared imagery). Furthermore, the algorithmic classification method does not include the deterministic bias of other studies in choosing the thermophysical class thresholds. However, it was not examined whether this provided an optimal partitioning of the dataspace, given that the maximum number of classes is chosen by the user. An optimal classification would have the class boundaries match the minima in pixel counts, and the number of classes would be approximately the number of maxima, so that the maximum amount of information is being extracted from thermal inertia and albedo. The correlation between the boundaries of the 10 thermophysical classes produced in this work and the density of points in global thermal inertia and albedo is visualised in Figure C-0-1.
302
Appendix C: Methodological Considerations for Chapter 8
Figure C-0-1: The partitioning of thermal inertia and albedo values into 10 classes from Chapter 8. The primary peaks occur in class 8 (pink), class 6 (light blue), class 4 (light green) and a broad peak that is spread across classes 2 and 3 (orange and yellow). Each class encompasses a broader range in thermal inertia than albedo, with the thermal inertia boundaries reflecting the underlying minima in pixels. Figure C-0-1 suggests that although most global maxima in the thermal inertia and albedo contour plot map to single classes (with the intervening classes incorporating intermediate values), one maximum is spread over two classes (class 2 and 3) indicating that those classes may be justifiably merged to form a larger class.
To assess how closely the boundaries of the 10 algorithmically
defined classes matched the underlying data, the density of points within each was examined through 3D histograms. The optimal classes (clusters) are identified by the counts (z) reaching a minimum at the boundary of the class in the two dimensions of thermal inertia (x) and albedo (y). This was done for each of the 10 classes, shown in Figure C-0-2-Figure C-0-11 below.
All classes showed a
minimum at their boundaries in the thermal inertia dimension. Classes 3 and 7-10 also showed a minimum at their boundary in both extremes of albedo, however classes 1-2 and 4-6 have an albedo minimum at only one extreme. This suggests that these two groups of classes could be merged in the albedo dimension,
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reducing the number of classes to 7. Figure C-0-5 reveals, however, that class 4 shows two local maxima in thermal inertia, suggesting that the class could be subdivided along the central thermal inertia local minimum. This behaviour is also observed clearly in classes 5 and 7. Hence if these classes were sub-divided 3 new classes would be defined. Two alternate partitions of thermal inertia and albedo are shown in Figure C-0-12. From this figure it is clear that stipulating a higher maximum number of classes maintains similar class boundaries in albedo but splits the classes to encompass smaller ranges in thermal inertia.
Given the large range in thermal inertia
observed at each landing site, reflecting a range from dust to bedrock [Squyres et al., 2004a; Vaughan et al., 2010], it is reasonable to incorporate heterogeneity in surface materials and hence a broad range in thermal inertia within each class. The arguments presented above indicate that although 10 classes may be around the optimal number of subdivisions in martian thermal inertia and albedo, the boundaries of the classes produced in this work could be adjusted to more closely align with the underlying data distribution. A finer partitioning in thermal inertia and albedo would also be useful, as it would provide greater discrimination of the degree of dust and particulates coverage of martian surfaces, and could potentially increase the agreement between the surface materials map and the geologic map (for example in Figure 8-13). Expanding the classification to a higher number of classes (such as 20) is therefore an avenue to explore in future work.
Alternate classification algorithms More sophisticated algorithms devised from artificial intelligence applications, such as fuzzy classifiers (including neural networks and other artificial learning algorithms [Mohanty, 1996]) and decision tree classifiers [Friedl and Brodley, 1997], have been developed in Earth remote sensing.
Neural networks in
particular have been shown to demonstrate a higher level of accuracy than ISODATA [Duda and Canty, 2002; Key et al., 1990] when applied to finding unsupervised clusters in multi-dimensional datasets [Bischof et al., 1992; Farrand et al., 2008; Perlovsky and Mcmanus, 1991; Ripley, 1994]. The main advantage of these algorithms over MAXLIKE is their treatment of mixels [Murthy et al., 2003], pixels that contain more than one distinct category of land cover (surface
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materials) and hence whose surface reflectance is a weighted average of contributions from each component [Lo and Choi, 2004; Zhang and Foody, 2001] (8.1.2). The MAXLIKE classifier and other hard classifiers employ Boolean logic to assign pixels to classes- each pixel belongs to one and only one class and must meet all the decision criteria for that class. Fuzzy classification is more flexible as it assumes that a class can be modelled as a continuous field rather than as an entity with a discrete boundary [Jiang and Eastman, 2000; Sicat et al., 2005]. This makes fuzzy classifiers more appropriate for handling continuous data as values that represent transition regions or attributes can belong to more than one class [Au et al., 2006].
Pixels assigned to a fuzzy class will each have a value
representing their degree of membership or suitability to the decision criteria. This relaxation of strict class membership allows for wider scope in selecting the members of a class and determining the suitability of the clustering [Baja et al., 2002]. Furthermore it is particularly useful in dealing with mixels, as pixels reflecting homogenous categories are separated and mixels are described as partial members of those pure categories, with the percentage membership reflecting the relative sub-pixel contribution of each [Zhang and Foody, 1998]. This provides a higher level of accuracy in the classification of mixels than a Boolean classifier [Binaghi et al., 1999; Foody, 2002], which would assign them to their highest probability class, reflecting either the dominant sub-pixel category or the average of the sub-pixel variation [Shanmugam et al., 2006]. The digital classification algorithms employed in this research (ISODATA and Maximum Likelihood) are well documented, easily accessible through commercial software, and have been demonstrated to be a sophisticated and rigorous method of uncovering the inherent statistical relationships between physical variables. Furthermore, they remove the classification bias inherent in previous derivations of thermophysical classes, as the algorithms are sensitive to the natural clustering within the datasets. Nevertheless, this above discussion provides motivation for future work with more sophisticated clustering algorithms to more accurately derive thermal inertia and albedo thresholds.
305
Figure C-0-2: Density of pixels in class 1 of the thermophysical map. Axes are: thermal inertia (x), albedo (y), count (z). Shading reflects cells that have > 1 pixel. The density of pixels falls to a minimum at high and low thermal inertia. It also shows a minimum at low albedo, but not at high albedo where it joins class 2.
Figure C-0-3: Density of pixels in class 2. The pixel density falls to a minimum at high and low thermal inertia. It also shows a minimum at high albedo where it joins class 3, but not at low albedo where it joins class 1.
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Figure C-0-4: Density of pixels in class 3. Class 3 falls to a minimum at high and low thermal inertia, and high and low albedo.
Figure C-0-5: Density of pixels in class 4. The pixel density falls to a minimum at high and low thermal inertia. It also shows a minimum at low albedo where it joins class 3, but not at high albedo where it joins class 5.
307
Figure C-0-6: Density of pixels in class 5. The pixel density falls to a minimum at high and low thermal inertia. It does not show a minimum at high or low albedo.
Figure C-0-7: Density of pixels in class 6. The pixel density falls to a minimum at high and low thermal inertia. It shows a minimum at high albedo where it joins class 7, but not at low albedo where it joins class 5.
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Figure C-0-8: Density of pixels in class 7. The pixel density falls to a minimum at high and low thermal inertia. It also shows a minimum at high and low albedo.
Figure C-0-9: Density of pixels in class 8. The pixel density falls to a minimum at high and low thermal inertia. It also shows a minimum at high and low albedo.
309
Figure C-0-10: Density of pixels in class 9. The pixel density falls to a minimum at high and low thermal inertia. It also shows a minimum at high and low albedo.
Figure C-0-11: Density of pixels in class 10. The pixel density falls to a minimum at high and low thermal inertia. It also shows a minimum at high and low albedo.
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Figure C-0-12: Scatterplot of all pixels showing the partitioning of the dataspace into 5 classes (above) and 15 classes (below). Thermal inertia x-axis is logarithmic. The discrete values of thermal inertia below 50 tiu are an artefact of the ‘lookup’ model used to derive thermal inertia (Equation 7-8).
311
GLOSSARY Term Absorption Active ADP Adsorption Aeolian Albedo Alpha particles Amazonian
Anaerobic Andesite Asthenosphere ATP Barophile Boiling
Brittle to ductile Chaotropic Chemosynthetic
Circumstellar habitable zone Classification (algorithmic/ unsupervised) Closure depth CMB Collegiative properties
Definition The permeation and dissolving of a fluid by a substance Organisms which are reproducing and metabolising at a measureable rate Adenosine diphosphate, a nuceleotide produced by ATP which is involved in energy transfer within cells The adhesion of a substance to a surface creating a film Processes driven by wind activity The fraction of incident solar energy in visual wavelengths that is reflected by the surface. Albedo of 1 indicates 100 % reflection Two protons and two neutrons bound together (a helium nucleus) Martian geologic epoch encompassing surfaces aged between 3.0Ga to the present. Late Amazonian encompasses surface ages between 400 Ma to the present Not requiring oxygen A grey volcanic rock dominated by plagioclase which is a major component of Earth’s crust (along with basalt) Ductile region of the mantle Adenosine-5'-triphosphate, an enzyme which transports chemical energy within cells for metabolism Pressure loving Condition occurring when the partial pressure Ppart of a substance (e.g. water) equals the surrounding atmospheric pressure Ptot at any temperature on the vapour curve Transition within the crust from rocks being prone to fracturing to rocks deforming by creep with the closure of fractures and pore space Substances that disrupt the intermolecular forces between water molecules and destroy complex molecules such as DNA and proteins Describes the biological process of converting carbon molecules into nutrients using oxidation processes to obtain energy rather than sunlight The range of distances from a central star where liquid water would be stable on a planetary surface A broad suite of algorithms that define multi-dimensional classes by employing: (i) some measure of similarity to define classes; and (ii) some formula for class membership to assign pixels Depth at which subsurface pore space is no longer in contact (or timescale of diffusive contact is sufficiently long) with the atmosphere Core-mantle boundary Properties of a substance that dependent only on the number of solute particles (mols) rather than the chemical identity of the constituents
312
Cratons CRISM Critical point
Deliquescence
Denaturation Dessiccated Dormant Duricrust
Dust cover index
Elastic lithosphere
Eutectic
Extremophile
Fines Gamma rays Geotherm GRS Habitable zone
Halophile HEND Hesperian
Glossary
Ancient, stable and thick regions of the continental crust that have not undergone deformation for ~ 104 Ma Compact Reconnaissance Imaging Spectrometer for Mars (onboard Mars Reconnaissance Orbiter) The termination of the vapourisation curve and the temperature at which the densities of liquid and vapour phase of a substance become equal A phase transition from a dry solid to a droplet of saturated solution, which occurs when a particle absorbs water from the atmosphere onto its crystalline surface until it is liquefied and dissolved The loss of structure in proteins and nucleic acids, which disrupts the cell activity of an organism Extremely dry with minimal or no water content Metabolism is unmeasurably low, no reproduction is occurring but basic cell repair may be underway Surface observed on Mars consistent of fine sand, 0.001-1.5 m grains, cemented together by salt or ice. The grains show a bulk thermal behaviour, rather than behaving as individual particles (fines) Parameter indicating the degree of dust coverage on the martian surface, derived from the average thermal emissivity at wavenumbers 1350-1400 cm-1. Low DCI indicates high dust coverage. The range of depths throughout which the crust displays brittle behaviour. Lower bound (ductile behaviour) is related to the thermal gradient within the crust and mantle
A particular mixture of substances with a single chemical composition that solidifies at a lower temperature than any other mixture of those substances. This composition is known as the eutectic composition and the minimum temperature is known as the eutectic temperature A microbial organism that thrives in conditions beyond the growth or tolerance limits of most organisms on Earth. Extreme conditions include temperature, pressure, salinity, UV radiation, etc. Individual grains not bonded to other grains, such as by a cementing agent (as in duricrust). Includes martian dust and sand High-frequency radiation with typical frequencies > 1019 Hz The geothermal gradient is the rate at which the Earth's temperature increases with depth, indicating heat flowing from the Earth Gamma Ray Spectrometer (onboard Mars Odyssey) A region where the environmental conditions meet the requirements of Earth life (for example, the temperature, availability of liquid water, nutrients, etc.). Could refer the circumstellar habitable zone or a region within a planet Salt loving High Energy neutron Detector (onboard Mars Odyssey) Martian geologic epoch encompassing surfaces aged between 3.7Ga to 3.0Ga
313
HiRISE Hydrostatic pressure Indurated Insolation ISODATA
JMARS Life space
Lithosphere Lithostatic pressure Lobate debris apron Maximum Likelihood
Metastable
Mixels Mohorovicic discontinuity (Moho) Noachian Nutrients
Obliquity OMEGA
Optical depth
High Resolution Imaging Science Experiment (onboard the Mars Reconnaissance Orbiter) Pressure corresponding to the overburden of water A hardening of the surface due to grains being cemented together (as in duricrust) A measure of solar radiation energy received on a surface Unsupervised clustering algorithm which iteratively passes through a dataset and defines clusters based on minimizing the pixel separation values Java Mission Planning and Analysis for Remote Sensing The full range of P-T conditions experienced by life on Earth. The related term ‘life box’ indicates the overall range of conditions that a given organism experiences
The rigid region of a planet (comprising the crust and outer mantle) where the rocks show brittle rather than ductile behaviour Pressure corresponding to the overburden of rock Smooth sloped cliffs with rock-debris at their base, most likely icy rock glaciers A supervised classifier which assigns pixels to classes based on the statistical description of the classes provided by an unsupervised classifier (e.g. ISODATA), and assuming that the probability of membership to each class is a multi-dimensional Gaussian A particular phase of a substance (solid, liquid, gas) is metastable when the total pressure and temperature, but not the partial pressure, are within the region of the phase diagram that corresponds to that phase being dominant. A metastable phase will have a long lifetime than an unstable one. Pixels that contain more than one distinct category and hence whose value is a weighted average of contributions from each component The density transition where seismic wave velocities increase with depth that marks the transition from crust to mantle Martian geologic epoch encompassing surfaces aged between 4.1Ga to 3.7Ga The elements which are used by life to provide the building blocks of biological structures, such as nucleic acids and proteins. Earth life uses C, H, N, O, P and S are used to build biomass, with carbon and nitrogen being the most critical Tilt of a planet’s axis Visible and Infrared Mineralogical Mapping Spectrometer / Observatoire pour la Minéralogie, l'Eau, les Glaces et l'Activité (onboard Mars Express) The logarithm of the fraction of light from the source that is not received at the detector, due to scattering or absorption by the
314
Osmophile Osmosis
Partial pressure (Ppart) Perchlorates Permeability Photolysis Potential martian biosphere (PMB) Probability distribution function Psychrophile Radiogenic Redox reaction
Regolith Rift zone Saltation Saturation SHARAD Skin depth SNC meteorite
Stable
Supercritical fluid
Glossary
medium; a measure of transparency Osmotic pressure loving The movement of molecules through a selectively permeable membrane, from low concentration to high concentration The pressure exerted by the molecules of the substance in the gaseous phase Salts consisting of a positive cation (eg. Na+, Mg2+) bonded with the perchlorate anion ClO4The connectedness of pore space which governs the flow of fluid. Material is ‘impermeable’ if its permeability is < 10-19 m2 The separation of intra-molecular bonds through the absorption of light All environments on Mars where the pressure and temperature are within the growth and metabolism thresholds of Earth organisms, and within the liquid region of the phase space of an martian liquid water The functional description of the likelihood of a variable having a certain value; normalised to 1, where 1 represents certainty Cold loving Produced through radioactive decay, the release of energy from an atom by the emission of ionising radiation. Indicates oxidation-reduction reactions. Oxidation reactions are those which increase the oxidation state or remove an electron form a substance. Reduction reactions are those which decrease the oxidation state or supply an electron. Layer of loose material covering solid rock, including dust, soil and broken rock Fissures on the side of volcanoes which allow the lateral release of lava (distinct from the vertical release of lava from the summit) The movement of particles through transient lift, by air or water, ending in transport back to the surface The maximum vapour pressure of water in the atmosphere, at a given temperature Mars Shallow Radar sounder (onboard Mars Reconnaissance Orbiter) Depth at which the temperature variations within the subsurface have decreased to 1/e of their surface value A group of meteorites which most likely originated from Mars and provide some young samples of the martian crust (ranging between 106-109 years old)
A particular phase of a substance (solid, liquid, gas) is stable when the partial pressure and temperature are within the region of the phase diagram that corresponds to that phase being dominant. A fluid at a pressure and temperature above its critical point and hence neither a fluid nor gas but shows some properties of both. Small changes in P-T result in a large change in density.
315
TES THEMIS Thermal inertia
Thermophile Thin film
Unstable
van’t Hoff factor
Water activity
Thermal Emission Spectrometer (on board Mars Global Surveyor) Thermal Emission Imaging System (onboard Mars Odyssey) The thermal inertia of a material is its ability to conduct and store heat and provides a measure of its resistance to changing temperature. Low thermal inertia materials rapidly heat and cool at their surface due to their poor ability to distribute heat through conduction Heat loving Unfrozen water adsorbed onto the surface of ice or grains, which can remain liquid at temperatures as low as –130 ˚C. The physical bonding between water and the adjacent surface is van der Waals forces A particular phase of a substance (solid, liquid, gas) is unstable when neither the total pressure or partial pressure are within the region of the phase diagram that corresponds to that phase being dominant. Accounts for the behaviour of a non-ideal solution, where the solute is highly volatile or strongly electrolytic (such as a saline solution). For an ideal solution the van’t Hoff factor is equal to the dissociation number of the solute (the number of ions in solution). For a non-ideal solution, the van’t Hoff factor is determined experimentally as the ratio of the colligative property in the electrolyte solute to the expected value for a non-electrolyte solution The equilibrium amount of water available for reactions, defined as the ratio of the water vapour pressure over the solution (Ppart) to the vapour pressure over pure water at the same temperature. Lower limit of 0.6 has been observed for life
316
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