Retrogressive Thaw Slumps on Mars: Constraints on Potential Formation Mechanisms

Home , Timekeeping on Mars

UNIVERSITY OF CALGARY

Jean-Michel Wan Bun Tseung

A THESIS

SUBMITTED TO THE FACULTY OF GRADUATE STUDIES

IN PARTIAL FULFILMENT OF THE REQUIREMENTS FOR THE

DEGREE OF MASTER OF SCIENCE

DEPARTMENT OF GEOGRAPHY

CALGARY, ALBERTA

APRIL, 2008

ISBN: 978-0-494-38308-7

Abstract

Subsurface temperature variations for the equatorial regions of Mars were obtained using a heat conduction equation in order to determine ground ice stability on Mars. The results indicate that under current summer conditions, subsurface temperatures above 0°C occur for a few hours at one metre deep. However, the upper 2 – 3 metres of the ground at these latitudes are thought to be desiccated. As such, the melting of ground ice through simple thermal conduction is unlikely to occur. Conversely, a reconstruction of the ground thermal regime for Mars at high obliquity suggests that changes in atmospheric pressure, water vapor content and average temperatures could result in a deeper penetration of the thaw front at mid- to high-latitudes. In these regions, the top of the ice-table should be at one metre deep or less and calculations indicate that a thaw front would reach the ice table and induce melting.

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Acknowledgements

My first and most sincere gratitude goes to my supervisor Dr. Brian J. Moorman for accepting me as one of his graduate students and for his perseverance in supporting me throughout my entire graduate program. His continued encouragement, patience, understanding and inspiration have greatly contributed in directing me not only towards the successful completion of my Master’s degree, but primarily and most importantly, towards becoming a better individual and friend. This program has been one of the most formative and exciting experiences of my life and would not have been so without his insight and knowledge. Thank you. I am grateful to my fellow friends and graduate students for fruitful discussions and for having provided me with advice, friendship and insight in various occasions. For that, thank you Pablo Wainstein, Christopher Stevens and Derek Wilson. My thanks also go to Chris Hugenholtz for his mentorship and friendship. I would like to acknowledge once more Dr. Brian Moorman and the members of my graduate proposal defence committee, Dr. Darren Sjogren and Dr. Yvonne Martin, as well as the members of my thesis defence committee, Dr. Rob Stewart and Dr. Darren Sjogren for their time, expertise and good- natured support in order to better my work. Thank you Paulina Medori for having supported me throughout my graduate program and for having kept an eye on the various administrative matters that were important in the completion of my degree. Thanks also to Brenda Mottle for her assistance and precious advices both as a lab technician and a friend. Thank you also Dr. Derald Smith and Dr. Andrea Freeman for your support and your trust in me during the two semesters I spent working with you. My thanks also go to the Department of Geography at the University of Calgary for the various scholarships and bursaries that helped me to give a good start to my graduate degree. To the Alberta Ingenuity Fund for their generous scholarship and offer for continued support without which I would not have been able to achieve what I did. Last but certainly not least, thanks to my wife for her unconditional support and patience in my life and for being the most wonderful partner on the ground and in the air.

Table of Contents

Approval Page...... ii Abstract...... iii Acknowledgements...... iv Table of Contents...... v List of Tables...... viii List of Figures and Illustrations...... ix List of Equations...... xi List of Symbols...... xii Epigraph...... xiii

CHAPTER ONE: INTRODUCTION...... 1 1 Introduction ...... 1 1.1 Rationale...... 1 1.2 Implications ...... 2 1.3 Hypothesis ...... 3 1.4 Aim and objectives...... 3

CHAPTER TWO: METHODS...... 5 2 Methods...... 5

CHAPTER THREE: LITERATURE REVIEW...... 8 3 Literature review...... 8 3.1 Permafrost...... 8 3.1.1 The freezing process...... 9 3.1.2. Moisture in frozen and unfrozen soils ...... 9 3.2. Underground ice...... 10 3.2.1. Ground ice...... 11 3.2.2. Buried ice...... 12

3.2.3. Ground ice or buried ice?...... 13 3.2.4. Ground ice conditions...... 13 3.3. Thermal properties of soils...... 15 3.3.1. Thermal conductivity...... 15 3.3.2. Heat capacity...... 16 3.3.3. Thermal diffusivity...... 17 3.4. The active layer...... 19 3.4.1. Active layer instability and detachment...... 20 3.5. Retrogressive thaw slumps...... 22 3.5.1. Formation and stabilization of RTS ...... 24 3.5.2. Ground ice content in retrogressive thaw slumps ...... 26 3.5.3. Material type and liquid limit...... 27 3.5.4. Size of RTS ...... 27 3.5.5. Retreat rates of RTS ...... 29 3.5.6. Summary of RTS on Earth...... 30

CHAPTER FOUR: MARS………………………………………………………………...31 4. Mars ...... 31 4.1. The Martian atmosphere and climate...... 33 4.2. The water cycle of Mars...... 36 4.3. Models of ground ice in the Martian crust ...... 38 4.4. Physical characteristics of the Martian fines ...... 44 4.5. Thermal properties of the Martian permafrost...... 46

CHAPTER 5: RESULTS AND DISCUSSION……………………………………………49 5. Results and discussion ...... 49 5.1. RTS formation under current climate ...... 50 5.1.1. Surface and subsurface temperatures ...... 51 5.1.2. Heat conduction results...... 57 5.1.2.1. Results based on published data ...... 58

5.1.2.2. Results based on raw MER data ...... 67 5.1.2.3. Thaw of ice-rich permafrost ...... 69 5.1.2.4. Melting of ground ice at temperatures below 0°C...... 76 5.1.3. Alternative RTS formation mechanisms under current conditions...... 77 5.2. RTS formation under past (warmer) climate...... 78 5.2.1. Orbital variations and climate change ...... 78 5.2.2. Melting of ground ice at high obliquity...... 81 5.3. Uncertainties...... 85 5.4. Sensitivity test...... 87 5.5. Retrogressive thaw slumps in Athabasca Valles...... 91 5.5.1. Features of the RTS in Athabasca Valles...... 93 5.6. Ground ice and RTS formation in Athabasca Valles...... 94

CHAPTER 6: SUMMARY AND CONCLUSIONS………………………………………98 6. Summary and conclusions ...... 98 6.1. Future research...... 100

REFERENCES...... 101

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List of Tables

Table 1. Thermal properties of soils and their constituents...... 19

Table 2a, b. Values of thermal and physical properties for Martian soils...... 48

Table 3. Comparison between published and calculated surface temperatures...... 86

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List of Figures and Illustrations

Figure 1. NASA’s theme “follow the water”...... 3

Figure 2. Methodology...... 5

Figure 3. Active layer detachment...... 21

Figure 4. Inactive retrogressive thaw slump...... 23

Figure 5. Ice face of a RTS and slump floor with levees and gullies...... 23

Figure 6. Partial coverage of an ice face with debris...... 25

Figure 7. Massive ice in a retrogressive thaw slump...... 26

Figure 8. Polycyclic retrogressive thaw slump...... 28

Figure 9. Possible pingos on Mars...... 32

Figure 10. Thermokarst landforms on Mars...... 33

Figure 11. Mars polar caps...... 35

Figure 12. Map of water equivalent on Mars...... 39

Figure 13. Ground ice model for Acidalia Planitia...... 41

Figure 14. Ground ice distribution on Mars...... 43

Figure 15. Martian surface at Opportunity’s landing site...... 46

Figure 16. Landing sites of Opportunity and Spirit...... 52

Figure 17a. Surface and near-surface temperatures at Spirit’s landing site...... 53

Figure 17b. Surface and near-surface temperatures at Opportunity’s landing site...... 54

Figure 18. Surface and subsurface temperatures for Meridiani Planum...... 59

Figure 19a. Annual temperature envelope for Meridiani Planum...... 60

Figure 19b. Temperature variations over one sol at Meridiani Planum...... 61

Figure 20. Hourly temperature variations for sol 137 of mission at Meridiani Planum.....62

Figure 21. Subsurface temperature profile for Meridiani on sol 137 of mission...... 63

Figure 22a. Annual temperature envelope for Gusev Crater...... 64

Figure 22b. Temperature variations over one sol at Gusev Crater...... 65

Figure 23. Hourly temperature variations for sol 1 of mission at Gusev Crater...... 66

Figure 24. Hourly temperature variations for sol 180 of mission at Gusev Crater...... 66

Figure 25. Annual temperature envelope for Meridiani...... 68

Figure 26. Annual temperature envelope for Gusev Crater...... 69

Figure 27a, b, c, d. Conceptual model of an ablating ice face on Mars...... 74

Figure 28. Martian obliquity cycle for the past five million years...... 80

Figure 29. Temperature envelope for Mars at high obliquity...... 83

Figure 30. Hourly surface and subsurface temperatures at high obliquity...... 84

Figure 31. Retrogressive thaw slump triggered by lateral stream erosion...... 85

Figure 32. Sensitivity test for the annual temperature envelope in Meridiani...... 88

Figure 33. Sensitivity test for the annual temperatures envelope in Gusev Crater...... 89

Figure 34. Sensitivity test for temperature variations over one sol at Meridiani...... 90

Figure 35. Sensitivity test for temperature variations over one sol at Gusev Crater...... 91

Figure 36. Subset of HiRISE image PSP-002793-1885...... 92

Figure 37. Subset of HiRISE image PSP-002437-1875...... 93

Figure 38. Inset of figure 36...... 94

Figure 39. RTS in N.W.T. Canada and HiRISE image PSP-002793-1885...... 96

Figure 40. Polycyclic retrogressive thaw slump in the Mackenzie delta...... 97

List of Equations

Equation 1. Calculation of active layer thickness...... 20

Equation 2a. Retreat rate of an ice face...... 29

Equation 2b. Retreat rate of an ice face...... 30

Equation 3. Amplitude of diurnal temperature variations...... 54

Equation 4. Subsurface temperature for a steady state condition...... 55

Equation 5. Amount of heat flow through conduction...... 55

Equation 6. Temperature at any depth z...... 56

Equation 7. Retreat rate of an ice face (same as 2b)...... 73

List of Symbols

Q* Net exchange of radiation between the Earth’s surface and the atmosphere QH Sensible heat QLE Latent heat QG Conduction of heat into the ground J Joules g Gram r Radius W Watts K Kelvin m Metre Kg Kilogram K Thermal conductivity k Thermal diffusivity C Volumetric heat capacity s Seconds V Retreat rate of an ice face F Energy flux juxtaposed on an ice surface (equation 3) Ice content T Difference between the temperature of ground from 0°C m Micrometre I Thermal inertia Density q Geothermal flux A Albedo 0 Porosity c Specific heat capacity Tz Temperature at depth z Ts Surface temperature z Depth P Period of a temperature wave in seconds As Amplitude of the surface temperature wave

Tz Mean surface temperature t Time in seconds

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“In the retroductive inferences of geomorphology, analogy serves merely to suggest fruitful working hypotheses, thereby leading to completely new theories that bind together any newly discovered facts. Mars’ landscape provides particularly stimulating opportunities to practise geomorphological reasoning, generating hypotheses that may initially strike some researchers as outrageous. Nevertheless, it is the productive pursuit of such hypotheses that leads ultimately to new understanding, not only of Mars, but of Earth itself” (Baker, 2001).

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CHAPTER ONE

1. Introduction

Since the early days of the Mariner and Viking missions, geomorphological evidence has shown that Mars, like Earth, is, or at least has been, a water-rich planet (Baker 1982). Important past roles of liquid water have been inferred from a plethora of landforms such as gullies, channels, alluvial fans, erosional valleys, salt-rich evaporite beds, and other flow features on Mars (Kargel 2004). However, the idea that Mars was once water-rich has proven difficult to conform to various characteristics of the Martian geochemistry and geophysics. Undoubtedly, theoretical premises have revealed controversial and include various atmospheric models for early Mars suggesting that: a) CO2 greenhouse would have generated warm and wet conditions (Pollack et al. 1987), b) refutation of the existence of such a greenhouse, and, c) that early in Mars’ history; high impact rates would have released water for valley formation. Nevertheless, what seems certain today is that the cryogenic state of Mars supports a greater role of solid H2O in the Martian crust compared to Earth’s crust, which is mostly thawed.

1.1. Rationale

In permafrost environments of Earth, ground ice features are diverse, representing the geomorphic expression of multiple processes that currently operate or have modified large parts of the Arctic and Antarctic regions through the formation, modification and degradation of ice-water features. The geomorphic characteristics of Mars suggest that its permafrost also should have gone through one or more periods when ice-water equilibria were important (Kargel 2004). This view is supported by geomorphic evidence characterized by landforms similar to terrestrial ground ice and periglacial features including pingos (Burr et al. 2005; Soare et al. 2005; Paige and Murray 2006), debris flows, frost mounds, polygonal patterned ground (Pechmann 1980; Seibert and Kargel 2001; Mangold 2005), thermokarst (e.g. Costard and Kargel 1995; Wan Bun Tseung et al. 2006),

2 protalus lobes (Mangold 2003), and rock glaciers (e.g. Colaprete and Jakosky 1998; Whalley and Azizi 2003). However, Retrogressive Thaw Slumps (RTS), one of the most common ground ice features found on Earth, have yet to be identified on Mars.

1.2. Implications

The now famous theme “follow the water” (fig. 1) adopted by NASA clearly demonstrates the importance of water in the search for life on extraterrestrial planetary bodies. Mars exhibits several lines of evidence that its surface was once carved by liquid

H2O. However studies indicate that under present climatic and atmospheric conditions, water (liquid or solid) cannot exist on the surface at the equator and lower mid-latitudes and would sublimate instantaneously if exposed to surface conditions. In contrast, theoretical modeling indicates that H2O ground ice could exist within the upper metres of the crust with its distribution varying latitudinaly. This assumption is supported by data including those generated by the neutron and gamma ray spectrometers on board the Mars Odyssey, which indicate high concentrations of hydrogen within the upper metre of the surface, interpreted as being water ice (Boynton et al. 2002; Feldman et al., 2002; Feldman et al. 2004b). Finding ground ice on Mars remains unlikely for present missions due to technological constraints and a lack of knowledge of the Martian permafrost. However, determining whether ground ice features such as RTS could exist on Mars would prove essential in planning and designing future exploration missions to Mars, including the selection of future landing sites. Furthermore, determining whether thaw features can occur on Mars would have important implications for past or present biotic activity on Mars.

Figure 1. NASA’s theme “follow the water” has been carried throughout the Mars Exploration Program since the year 2000. Source: NASA.

1.3. Hypothesis

This research looks at the variability in environmental conditions between Earth and Mars in order to determine whether retrogressive thaw slumps can or could have formed on Mars.

1.4. Aim and objectives

While a number of studies have attempted to explain landforms on Mars using terrestrial permafrost/periglacial analogs and processes to fit Martian features, few have investigated terrestrial landforms and their possible existence on the Martian surface using a process-based approach (i.e. determine whether Earth-based processes can or could have occurred on Mars instead of attempting to fit terrestrial processes into Martian landforms). Furthermore, a thorough understanding of Earth-based processes is crucial if one is to understand similar mechanisms on Mars.

The purpose of this study is to investigate the possibility of RTS formation on Mars primarily through the analysis of heat conduction properties for the southern equatorial regions of Mars while taking into account the variability between Earth’s and Mars’ ground thermal and physical properties. The objectives are to evaluate the evidence from Mars, consider alternative models and draw conclusions regarding whether Martian conditions allow or could have allowed the formation of RTS on its surface.

CHAPTER 2

2. Methods

This chapter outlines the methods used in this study to obtain results for the formation of retrogressive thaw slumps on Mars. Figure 2 is a graphical outline of the methodology undertaken in this thesis.

Figure 2. Outline of methodology.

In order to answer the research question, a literature review of Earth-based permafrost, particularly RTS, was accomplished in order to constrain the parameters involved in the triggering and formation of RTS as well as the factors that affect it. The

6 literature review primarily included published literature but was also supported by the examination of aerial photographs and the quantification of RTS morphology. Similar information was then obtained for Mars (when possible) and included details on the physical and thermal characteristics of the Martian soil. Because the occurrence of ground ice on Mars is subject to disparate views, several published models were taken into consideration and a depth to ice table was then chosen for this study. This depth was based on existing models, theoretical premises and data including gamma ray and neutron spectrometry data as well as quantitative and qualitative geomorphic evidence. Because not all information available for Earth can be obtained for Mars, some generalisations had to be made such as assuming a uniform particle size distribution to depth. All these are presented in more details in a later section. Without the use of instruments for on-site analysis or data recordings, different mathematical and/or physical models are used to investigate the thermal characteristics of a ground thermal regime such as thermal inertia and gamma ray spectrometry in the case of Mars. However, these models only provide general estimates for subsurface temperatures. Several approaches could be used to answer the research question of this thesis. However, determining subsurface temperatures in the shallow subsurface (upper metres) appears the most appropriate method since the formation of retrogressive thaw slumps is largely dependent on a disturbance in the temperature regime of the ground. As such, a heat conduction equation using surface temperatures is used in this study. The heat conduction equation used in this study allows the calculation of temperatures at any depth z, essential to determine the depth of penetration of a thaw front in to the ground. Details of the heat conduction equation are discussed in section 5.1.1. To account for variabilities in the results, data (surface temperatures) in ASCII format for Mars were acquired from two sources: 1. the published literature (Spanovich et al. 2006), and, 2. the Planetary Data System (PDS), repository of most missions’ data to Mars. The PDS data was then processed to obtain a data format (temperature values) usable in the heat conduction equation. The surface temperatures obtained were then computed using the heat conduction equation in order to reconstruct the thermal regime of the Martian permafrost under current conditions at two locations, Meridiani Planum and Gusev

Crater. For the current climate, results obtained from the two datasets were then compared. This study uses the heat conduction equation (6) provided by Williams and Smith (1989) that allows the calculation of subsurface temperatures accurately for the upper metres of the ground. Similar calculations were then accomplished for a past hypothetical climate using surface temperatures modelled for the last obliquity cycle (100,000 years ago), also obtained from the published literature.

CHAPTER 3

3. Literature review

The discipline of periglacial geomorphology emerged in the 1890s. Polish geologist, Walery Lozinski, delivered the first unifying concept of periglacial geomorphology (Jahn 1954; French 2000a) by putting forward the concept of a periglacial zone that encompasses environments adjacent to glaciers (Lozinski 1909; French 2003). This concept was later altered, primarily in order to differentiate between cold glacial and cold non-glacial landscapes. Thus, the modern usage of the term “periglacial” refers to a wide range of cold, non-glacial environments where freeze and thaw (also referred to as frost action) and permafrost processes dominate, regardless of geographical location (e.g. Jahn 1975; French 1976a; Washburn 1979; French 1996). As such, two distinguishing criteria, one not necessarily exclusive of the other, are recognized in periglacial environments. The first is represented by the freezing and thawing of the ground, often associated with the presence of water. As concluded by J. Tricart., ‘the periglacial morphogenetic milieu is that where the influence of freeze-thaw oscillations is dominant’ (Tricart 1968, p830). The second diagnostic feature of periglacial environments is the presence of perennially frozen ground or permafrost. According to Péwé, ‘permafrost is the common denominator of the periglacial environment, and is practically ubiquitous in the active periglacial zone’ (Péwé 1969, p4). As such, he considers all regions currently underlain by permafrost to represent the modern periglacial zone.

3.1. Permafrost

Permafrost is widely defined as ground that remains below 0°C for two or more years (e.g. Washburn 1979; French 1996; van Everdingen 2002). Permafrost conditions in the western Canadian Arctic result from the complex pattern of glaciations during the Pleistocene and Holocene eras with permafrost of Pleistocene age is generally much thicker than in more contemporary conditions (Rampton 1982). Because the definition of

9 permafrost is purely thermal, permafrost therefore bypasses the issue of whether or not the presence of ground water is in the liquid or solid (frozen) state (Thorn 1992).

3.1.1. The freezing process

The freezing process is complex and its details are still not fully understood. While some frost action processes remain to be explained, it is clear however that different soils freeze at different rates depending upon the thermal properties and moisture content of the soil. In many instances, soils do not freeze when temperatures fall below 0°C but exist in a supercooled state. One common condition for the latter state is when saline ground water lowers the temperature at which soils freezes. Furthermore, the rate and amount of freezing will also depend on the duration and intensity of a temperature drop below 0°C (French 1996). The significance of annual freezing largely depends on the depth to which winter cold waves penetrate and the nature of the materials. It has been shown that the penetration depth of frost can vary by a factor of two or more even on lateral distance of metres (Williams and Smith 1989). Generally, the controlling factors on frost penetration depth are: a) the mean annual ground temperature, b) the amplitude of the annual surface temperature cycle, c) the nature of the ground surface cover, particularly snow cover, and, d) the thermal properties of the soil. At the beginning of the cold period, freezing advances at a slow rate downward from the surface but then accelerates dramatically at depth. This results from: a) upward freezing of the ground from the permafrost beneath (Mackay 1973c, 1974a), and, b) the decreasing moisture content with depth since soil water is initially drawn upwards to the freezing front, thus favoring the effects of latent heat release in the upper surface layers (French 1996).

3.1.2. Moisture in frozen and unfrozen soils

The presence of moisture in the active layer complicates our understanding of active layer processes since not all water freezes at 0°C. The active layer is the layer that freezes and thaws every year and is discussed in more details in section 3.5. In some frost

10 susceptible soils, as much as 40% of the water content may remain unfrozen at -1°C (e.g. Williams 1976). A further difficulty is the differentiation between the seasonally thawed layer possessing a temperature < 0°C (i.e. seasonally cryotic) and the uppermost layer of permafrost which is affected by seasonal thaw but remains below 0°C (i.e. seasonally active permafrost). French (1996) notes that it is also useful to distinguish between the freezing front, which represents the margin between frozen and unfrozen soil, and the cryofront, which is the 0°C boundary, when investigating active layer freezing during the fall and winter. Liquid water in frozen ground moves in the direction of ground temperature decrease in response to an imposed thermal gradient, a process commonly referred to as cryosuction. As explained by Williams and Smith (1989), when a freezing front advances from the surface downwards into the soil profile, ice forms in the pores of the soil matrix, resulting in the confinement of the unfrozen pore water. This decrease in the amount of water is further affected by adsorption and capillarity, the latter being the cause of increased molecular forces at the pore interface. Similar to the rise of water in a capillary tube due to the confinement of the meniscus between air and water, ice formation in a freezing soil results in water being confined to smaller spaces. This in turn increases the suction effect caused by the freezing soil and is responsible for the migration of water to the freezing zone. This has significant implications for the ice distribution when the active layer freezes. As a result, the ice content in the upper part of the active layer will increase when ground temperature is warmer than air temperature in winter while the lower parts of the active layer will witness an increase in ice content during the summer due to colder temperatures originating from the permafrost below.

3.2. Underground ice

While it is well known that the unique features found in cold regions are often the result of freeze and thaw processes, it is important to note that in most instances the presence of ice in the subsurface plays a major role in determining the final characteristics of those landforms. On the other hand, subsurface ice can be categorized into two types, ground ice and buried ice, according to their primary origin. For the last two decades, the

11 genetic classification of ground ice by Mackay (1972a), which revised previous Soviet classifications (e.g. Shumskii and Vtyurin 1966), has acquired overall acceptance within the scientific community in North America. This classification is based upon the water source prior to freezing and the principal transfer process, which moves water to the freezing plane. The major difference between Mackay’s and Soviet classifications is that the former did not include buried surface ice.

3.2.1. Ground ice

According to Mackay’s (1972a) classification, ground ice refers to ice in the substratum that is not of glacial origin. As such ground ice can be found in pores, cavities, voids and other openings in soil or rock and includes massive ice (ACGR 1988, p46). Because it encompasses various forms of frozen moisture, ground ice is common in many areas underlain by permafrost and varies greatly in depth, amount and type. In modern periglacial literature, ground ice may be epigenetic or syngenetic. Epigenetic refers to ice that develops inside the enclosing rock or soil after the latter has formed. Syngenetic ice refers to ice that forms at, or almost at, the same time as the parent sediment is being deposited and is thus often related to surface aggradation (French 1996). Ground ice content in permafrost can vary from imperceptible in some igneous and metamorphic rocks to extensive in unconsolidated, fine-grained Quaternary-age rocks. Field data acquired by Parmuzina (1978), Cheng (1983), and Mackay (1983) showed that important changes in ice content in the active layer occur in the winter months. In summer, a negative ground temperature gradient will result in liquid water migrating downwards and an increase in ice content in the upper part of the permafrost (Mackay 1983; French 1996). These studies correlate with well-known evidence that the highest ground ice contents occur in the top 1 – 5 m of permafrost and that an ice-rich zone is typically found between the base of the active layer and the top of the perennially frozen permafrost (e.g. Vtyurina 1974; Ershov 1979; Ershov et al. 1980; Mackay 1981; Shur 1988). Additionally, the upward and downward movements of unfrozen water in the active layer differ from one year to the next because the upward migration of unfrozen pore water occurs in winter (due to an impose temperature gradient; colder at the surface) when the

12 unfrozen pore water content is low. Conversely, a negative ground temperature gradient in summer (colder at depth due to the presence of permafrost) and a higher pore water content due to warm temperatures in the thawed layer result in a net increase in water migration downward to the permafrost compared to the upward migration during freezing (French 1996).

3.2.2. Buried ice

In contrast to ground ice, buried ice refers to ice of external origin. Several sources may account for such ice and include buried glacier ice from a past glaciation period, buried snow, buried lake and river ice, and, buried icings. As documented by Williams and Smith (1989), lake or river ice is probably the most common form of buried ice due to the inherent instable nature of many shorelines which result in the collapse of sediments onto the ice. On the other hand, while buried glacier may be less common than lake or river ice, the former may reach a much larger size than the latter due to the fact the entire sections of glaciers may become buried under mass wasting events, especially in mountainous regions prone to landslides. As a result, buried glacier ice has the potential to reach several kilometres in size. Buried glacier ice may also take the form of an ice-cored moraine where debris falls onto the lateral edges of a glacier and protects the ice from melting (Moorman and Michel 2000). Although reports of buried glacier ice are common in the former U.S.S.R., they tend to be fewer in Canada (e.g. Lorrain and Demeur 1985; Dallimore and Wolfe 1988; French and Harry 1988, 1990). Icings refer to accumulations of ice on the ground surface or in rivers and are usually fed by subsurface water. Their size may range from metres to hundreds of metres in diameter and they can be metres to tens of metres thick (Williams and Smith 1989; Moorman and Michel 2000). Several types of icings are recognized and include ground icings, spring icings and river icings. Ground icings form where water saturated ground presents little relief or no permanent surface drainage path. In the case of spring icings, a water spring gradually builds successive layers of ice and in some instances; this process persists throughout the winter freezing if the temperature of the spring water is slightly above the freezing point. River icings form when there is an accumulation of overflowing

13 water that freezes on top of the original surface ice during winter seasons (Williams and Smith 1989). While icings normally refer to surface ice that may subsequently be buried by landslides or other material displacement, they can also form below the surface. The process is similar to that of spring icing formation except that in the case of underground icing, the water being forced upwards in unfrozen soils may encounter a progressively downwards movement of seasonal freezing. In such cases, the spring water may freeze within the pores if the latter are coarse enough so that the remaining water is constrained between the underlying permafrost and the overlying seasonally frozen ground, causing a blister (a temporary mound or frost warp) (Williams and Smith 1989).

3.2.3. Ground ice or buried ice?

It is now widely accepted, at least in the north-American literature, that ground ice and buried ice have contrasting origins (e.g. Mackay 1972a, Williams and Smith 1989; French 1996). A recent trend, however, has seen North-American scientists include buried ice as ground ice. For simplicity, this study will consider both ground ice formed epigenetically or syngenetically and buried ice as ground ice. Furthermore, field studies have shown that determining the origin of underground ice is extremely complex, leading to various formation hypotheses. Such discrepancy is reflected in one of the most famous examples of exposed massive ice found at Peninsula Point along the Beaufort sea coast which is interpreted as some as being intrusive ice from glacial meltwater (e.g. Mackay and Dallimore 1992), buried snowbank (Fujino et al. 1983, 1988) and intrusive ice with multiple sources (Moorman et al. 1998). Furthermore, in retrogressive thaw slumps, it is the amount of ice (regardless of its origin) and the energy available to melt the ice that will ultimately control the growth or stoppage of growth of the slump.

3.2.4. Ground ice conditions

Two quantitative parameters are used to describe ground ice conditions. The first one is the ice content, which refers to the weight of ice to dry soil expressed as a

14 percentage. Soils having ice contents < 40 – 50% are generally referred to as low ice content soils whereas soils (usually fine-grained) with 50 – 150% ice content are considered as soils with high ice contents (French 1996). The second parameter is the amount of excess ice. Excess ice is defined as ice that exceeds the total pore volume available to it. It is that portion that exceeds saturation that is termed excess ice and can become a substantial source of water, also referred to as supernatant water, upon thawing of such an ice-rich ground. The volume of supernatant water in a thawed column of frozen soil is usually expressed as a percentage of the total volume of sediment and water. For example, relative volumes of supersaturated sediments (300 cm) and supernatant water (500 cm) would represent an excess ice value of 62.5 %. Although ice lenses may suggest supersaturated soils, upon thawing the water released by the ice lenses may not be enough to fill up and exceed pore space available. Soils containing excess ice are often termed ice-rich sediments/soils and may reach values as large as 70 - 80 %. Excess ice can materialize in various forms including segregated ice, buried ice, ice wedges and massive tabular ice (Permafrost Subcommittee, Associate committee on Geotechnical Research 1988). Massive ice refers to a body of pure ice that constitutes at least 250% on an ice-to-dry-soil weight basis (Mackay 1971). Massive ice can be in the form of ice-wedges, buried ice (surface or glacier ice) and intrasedimental ice (grown in place by the freezing of water within a sedimentary sequence), which includes intrusion ice, segregation-intrusive ice and segregation ice (Mackay and Dallimore 1992; French 1996). In the Arctic regions of Canada, Alaska and Russia, the melting of massive ice often leads to thermokarst and/or RTS formation. Although studies of ice content to depth are limited, most massive ice tends to occur within 50 or 60 metres of the ground surface with the greater concentration near the surface (Williams and Smith 1989). In some instance, however, massive ice has been encountered at depths exceeding 30m and sometimes 100- 200m (Mackay 1973b; Dubikov 1982).

3.3. Thermal properties of soils

An understanding of the thermal properties of soils is essential when interpreting and analyzing the growth or decay of ground ice features. This premise is reflected in the fact that the presence of water or ice close to their transition temperature produces a significant effect on the thermal properties of frozen soils and subsequently, ground ice forms. As stated by Gold and Lachenbruch (1973), the change of water from its liquid to solid form increases its conductivity four-fold, decreases its heat capacity by half, and releases heat equivalent to that required to raise the temperature of an equal volume of rock by about 150°C. This process, however, becomes more complex since the amount of unfrozen water and ice within the substratum is temperature dependent, resulting in considerable changes in soil thermal properties even with slight variations in temperature below 0°C. Furthermore, the inherent changing nature of soil conditions, sometimes within centimetres, such as soil type, density water content and temperature hinders a unique classification of thermal properties for entire soils profiles or formation (Williams and Smith 1989). As a result, even though laboratory experimentation can shed light upon these parameters, estimates and approximations must sometimes be used when studying ground ice features, especially when specific parameters such as soil type, density and water content are not known. Table 1 below provides the thermal properties of soils and their constituents.

3.3.1. Thermal conductivity

The thermal conductivity of a soil refers to the bulk property of that material that indicates its ability to conduct heat. However, because ice and water content within a soil are temperature-dependent, there exists no specific value for the thermal conductivity of a frozen soil that contains water. As a result, studies pertaining to the thermal conductivity of frozen soils using laboratory experiments cite specific soil types with known amounts of both ice and water. As shown in table 1, soil mineralogy is a determining factor when considering the thermal conductivity of a material. The conductivity of ice (2.24 W m-1 K-1) is almost five times greater than that of water (0.56 W m-1 K-1) but four times less that

16 quartz-bearing soils such as silt. Hence, for example, a silty soil in which water freezes at about -1°C will exhibit a significant and abrupt increase in thermal conductivity as the water passes from its liquid to solid form. In the case of sand whose most common constituent in inland continental settings is quartz, the thermal conductivity will be high. However, sands are known to be poor aquitards (compared to silt or clay) and retain little water, especially when topography promotes the movement of water under gravitational force (i.e. down slope).

3.3.2. Heat capacity

Heat capacity is a general term that represents the ability of a body to store heat as it changes in temperature and is measured in joules per Kelvin. However, because heat capacity is influenced by the size of the object (e.g. an ocean will have a larger heat capacity than a lake) it is often more useful to divide the heat capacity by the body’s mass in order to obtain a specific heat capacity (also called mass heat capacity or specific heat). Thus, specific heat capacity ceases to depend on the size of the material but instead depends on the material type (i.e. silt, clay, ice, etc). Specific heat capacity therefore describes the amount of heat required to change the temperature of 1 kg of material by 1 Kelvin (Williams and Smith 1989). It is also known that for a given amount of heat injected into a substance, the lower the heat capacity, the greater the temperature changes will be in that substance. As indicated in Table 1, the primary mineral constituents of most soils (quartz, feldspar and clay minerals) have similar mass heat capacity values compared to organic material, which has a higher specific heat, but because the latter is less dense it possesses a similar volumetric heat capacity (mass heat capacity multiplied by density of material) than the mineral constituents. As such, it is well known that the heat capacity of most dry soils range between ~1000 – 1500 J m-3 K-1 near 0°C. Of great importance however, is the addition of water (which has a specific heat of about almost 4200 J kg-1 K-1, nearly four times that of dry soils) to a soil; the specific heat of that soil will increase significantly and linearly as water content increase. On the other hand, dealing with water in frozen soils (which is temperature dependent) implies that small changes in temperature may greatly

17 affect heat storage or release through the form of latent heat. This is particularly true within a few degrees below 0°C. Latent heat refers to the amount of heat released or absorbed by a substance during a change of phase. Two types of latent heat are commonly recognized, the latent heat of fusion and the latent heat of vaporization. Each indicates the direction of heat flow from one phase to the next. As such, the latent heat of fusion represents the amount of heat that must be absorbed by 1 gram of a given substance in order for it to change from a solid to liquid state, and also marks the temperature at which melting occurs. In contrast, the latent heat of vaporization characterizes the amount of heat required to change a substance from its liquid to its gas form, the transition being commonly known as the boiling point. While the latent heat of vaporization in frozen soils is extraneous due to the cold temperatures dominating such regions, the latent heat of fusion can have a great effect on the unfrozen water content in a frozen soil. This latent heat effect is commonly referred to as the apparent heat capacity of the soil because it combines both the specific heat capacity of that soil and the heat released or absorbed (latent heat) during a phase change. In the case of water, the value of latent heat release corresponds to 333 joules per gram as mentioned above. In the case of frozen soils, it is when water changes from liquid to solid that the latent heat plays a great role. Following the example given by Williams and Smith (1989), considering 1 m3 of moist clay soil at 1°C and a heat capacity of 3 MJ m-3 K-1, one would require the removal of 3 MJ in order to lower the temperature to 0°C. If cooling continues, latent heat will be released and between 0°C - -1°C, this would amount to ~63.5 MJ m-3 for the moist clay soil.

3.3.3. Thermal diffusivity

Thermal diffusivity is the ability of a substance to conduct heat as the surrounding temperature changes and is measured in W m-1 K-1. As such, a substance with a high thermal diffusivity means that it will conduct heat rapidly and will quickly adjust its temperature to that of its surrounding.

Density Heat Thermal Thermal (kg m-3) Capacity Conductivity Diffusivity

(J kg-1 K-1) (W m-1 K-1) (x 10-6 m2 s-1) Soil constituents Quartz 2660 800 8.80 4.14 Clay minerals 2650 900 2.92 1.22 Organic matter 1300 1920 0.25 0.10 Water (0°C) 1000 4180 0.56 0.13 Ice (0°C) 917 2100 2.24 1.16 Air 1.2 1010 0.025 20.63 Water Unfrozen soils content (m3 m-3) Sandy soil (40% 0.0 1600 800 0.30 0.24 porosity) 0.2 1800 1180 1.80 0.85 0.4 2000 1480 2.20 0.74 Clay soil (40% 0.0 1600 890 0.25 0.18 porosity) 0.2 1800 1250 1.18 0.53 0.4 2000 1550 1.58 0.51 Peat soil (80% 0.0 300 1920 0.06 0.10 porosity) 0.4 700 3300 0.29 0.13 0.8 1100 3650 0.50 0.12

Table 1. Thermal properties of soils and their constituents (after Williams and Smith, 1989).

However, because thermal diffusivity is the ratio of thermal conductivity (K) to volumetric heat capacity (C), its variation with water content and temperature depends on

19 the interaction of the two latter properties. As documented by Williams and Smith (1989), the thermal diffusivity of frozen soils is highly dependent on temperature and primarily controlled by the heat capacity, especially between 0°C and -3°C. For example, as temperature decreases, the conductivity of a moist clay soil increases due to the formation of ice (which has a much higher conductivity than water) but the specific heat capacity decreases as the water turns from liquid to solid.

3.4. The active layer

The thermal profile of permafrost comprises an upper thin layer (ranging from centimetres to more than one metre thick depending on the geographic location) called the active layer. The active layer refers to the layer that thaws in summer and freezes again in winter (Muller 1947). As such, it is thinnest in polar regions (as little as 15 cm) and thickens towards sub-arctic regions sometimes exceeding 1 m in thickness. According to French (1996), the active layer reaches the top of the permafrost table in the zone of continuous permafrost whereas in discontinuous permafrost zones, the active layer may be separated from the permafrost by a zone of unfrozen ground called talik or a residual thaw layer. Because the thickness of the active layer is dependent upon ambient air temperature, slope aspect, vegetation, drainage, snow cover, soil and/or rock type and water content, it may vary from year to year depending upon the interaction of the controlling factors. Furthermore, since the active layer includes all ground that thaws and freezes every year, it encompasses the uppermost part of the permafrost that has a salinity or clay content that allows the permafrost to thaw and refreeze each year. This is possible as long as the material remains cryotic (below 0° C) (French 1996). The penetration depth of a freezing front is greatly dependent upon the intensity of the cold wave, its duration, and, the thermal and physical properties of the soil and rock, the overlying vegetation and snow cover. Knowing the values of thermal conductivity for frozen soil and latent heat of water in the soil, it is possible to calculate the depth of seasonal freezing and thawing (Washburn 1979). Gold and Lachenbruch (1973) give an estimate of the active layer thickness, x (cm), with the following:

= / log ()/TAePx 00 (1)

Where is the soil thermal diffusivity, P is the period of the temperature cycle, A0 is the surface temperature amplitude (the maximum displacement from a zero value during one period of an oscillation) and T0 is the mean annual surface temperature. Johnston (1981) and Williams and Smith (1989) give typical values of soil diffusivity (m2/s).

3.4.1. Active layer instability and detachment

One of the most significant causes of active layer instability relates to the time- dependent compression of the ground which is the result of ground thawing and drainage of pore water, commonly known as thaw consolidation (Morgenstern and Nixon 1971). In coarse-grained soils, water released by the thawing ground is unobstructed and the settlement of material is uniquely controlled by the position of the thawing front. In contrast, when thaw occurs in fine-grained soils, flow is restrained and the rate of settlement is controlled by the permeability and compressibility of the thawed ground. In such cases, a high rate of thaw may result in water released at a faster rate than it can flow from the soil, resulting in higher pore water pressures and subsequently, instability in slopes. As documented by Morgenstern and Nixon (1971), substantial pore pressure at the thawing front is likely to result in slope instability due to a decrease in shear strength at that plane. In many permafrost regions, the permafrost table serves as a lubricated slip plane and usually controls the depth of the failure plane. Such failures are commonly known as active layer detachments (e.g. Hughes 1972a; French 1996). Active layer detachments depend upon: a) soil moisture saturation following precipitation, or, b) high pore water pressures resulting from exceptionally rapid and deep thaw, both of which result in the detachment of the active layer from the underlying permafrost. In some cases, such failures can take the form of a mudflow whereas in others, a distinct slump scar is formed (fig 3.). While active layer detachments are common in the Arctic regions, they are difficult to predict (e.g. Hodgson 1982; Stangl et al. 1982;

Lewkowicz 1990). In situations where significant bodies of ice are exposed in the slump scar, retrogressive thaw slumps are often triggered.

Figure 3. Active layer detachment in the Mackenzie delta. Image credit: Brian Moorman.

One of the most important aspects to note is that active layer detachments can occur on low angled slopes. Although many failures occur on slopes of less than 15°, some have been recorded on slopes as low as 3° and are explained by the role of excess pore-water pressure created by the thawing ice. This phenomenon is typical of areas where during summer thaw, meltwater is produced more quickly that it can escape or drain from the soil thereby increasing the pressure and increasingly supporting the weight of the overlying material. As a result, total loss of friction may occur and the active layer slides over a layer of water (McRoberts and Morgenstern 1974b; Dyke 2000).

3.5. Retrogressive thaw slumps

Considered by some scientists as a type of back wearing thermokarst, RTS represent one of the most rapid erosive processes operating in contemporary periglacial environments. Thermokarst features are the result of the thawing of ice-rich permafrost due to a thermal disequilibrium in the thermal regime of the permafrost. This process results in the melting of ground ice and the subsidence of surficial material, forming a water-fill depression or basin called thermokarst lake (French 1996). Like thermokarst, retrogressive thaw slumps result from the thaw of ice-rich permafrost. Retrogressive thaw slumps reach their largest dimensions and most frequent occurrence in areas underlain by massive ground ice or unconsolidated ice-rich sediments (French, 1996). The southern Beaufort Sea region of the Canadian Arctic represents one of the most ice-rich permafrost areas with extensive massive ground ice and numerous RTS (Pollard and French 1980; Pollard 1990; Lantuit 2004).Morphologically, retrogressive thaw slumps (fig. 4) imitate earth flows common to more temperate regions but do not exhibit characteristics typical of rotational slides due to the lack of surface rupture. Retrogressive thaw slumps are often hundreds of metres in diameter but can be larger than a kilometre and commonly exhibit: a) a vertical headwall comprising the active layer underlain by a layer of ice-poor material, b) an ice-rich face (fig. 5) with an angle that varies between 20 and 50° (having either excess or massive ice) that degrades by ablation resulting from sensible heat fluxes and solar radiation as well as latent heat release from melting ice, c) a slump floor at the base of the headwall on which mud pools are often present due to the slumping and flow of thawed material, and, d) a slump lobe commonly characterized by gullies or mud levees (fig. 5). Where hummocks are present on the surface, minor cracks in the active layer are created as the hummocks detach and fall down the face (Lewkowicz 1987; de Krom 1990; French 1996).

Figure 4. Inactive retrogressive thaw slump. Stabilization can be inferred by the presence of vegetation on the slump floor and the complete coverage of the ice face. Image credit: Brian Moorman.

Figure 5. Ice face with massive tabular ice exposed. Gullies and levees can be observed on the slump floor channeling debris flows. Image credit: Brian Moorman.

3.5.1. Formation and stabilization of RTS

Retrogressive thaw slumps are commonly triggered following the removal of the layer of material that thermally protects massive or excess ground ice, although other mechanisms may produce similar results. Mackay (1966) identified three main processes involved in the formation of retrogressive thaw slumps: a) detachment and sliding of the active layer or erosion of surface materials, b) mechanical or thermal erosion of the exposed ice, and, c) slumping or collapse of surficial material from above the frozen scarp face. The latter process is considered to be the most significant in terms of material removal and involves the gradual thawing of the scarp face as the ice is exposed (Mackay 1966). On eastern Banks Island numerous slumps occur on steep slopes and have been initiated by lateral stream erosion or wave action undercutting the base of the slopes. Other RTS are located alongside lakes or the coast where wave action or ice pushing may trigger RTS formation. In many areas, RTS activity is proportional to the intensity of wave action and coastal activity (de Krom 1990). Elsewhere, slumps are inferred to have been triggered by active layer slides. Such phenomenon has been widely reported in the high Arctic regions (e.g. Bird 1967; Lewkowicz 1992b) where over 50% of the total annual precipitation falls as summer rain resulting in the development of high pore water pressures. Similar conditions are also common during years of exceptionally high winter snowfall and/or rapid and late thawing. The stabilization of a RTS is a function of the balance between the amount and rate of debris supply and removal from the base of the headwall. This, in turn, is dependent upon the relative height of the ice-rich sediments in relation to the base of the headwall and the slope angle. Sediment accumulation at the base of the headwall resulting from material collapse helps in covering part of the exposed ice (fig. 6.), thus reducing the surface area exposed to solar radiation. Over time, complete covering of the ice body results in the stabilization of the RTS (de Krom 1990). This process is in great part controlled by the efficiency of flowage (drier material will tend to aggregate while material with a higher water content will flow and dissipate) of the slumped sediment and will determine the amount of material that will accumulate at the base of the scarp (resulting in ice coverage) or flow into the ocean or a river thus preventing

25 a layer of protective sediment (or its thickness) to accumulate at the base of the headwall. Ultimately however, the thickness of material required to protect an ice face from further melting will depend upon the depth of the seasonal active layer. It is this factor that will determine whether the summer thermal wave will penetrate deep enough to reach the ice or not. This active layer thickness will in turn depend on the intensity and duration of the summer heat waves as well as vegetative cover if the latter is present. Common values for the stabilization process often spread over periods ranging between 5 – 10 years for a RTS ~ 100 m in diameter (width).

Figure 6. Partial coverage on an ice face (note the presence of massive ice) by slumping of the material above. Image credit: Brian Moorman.

3.5.2. Ground ice content in retrogressive thaw slumps

Because retrogressive thaw slumps form due to the melting of ground ice, it is this ground ice content that will determine the geometry of a retrogressive thaw slump. Unlike thermokarst where excess ice represents the minimum amount of ice required in order for thermokarst degradation to occur, ice content in RTSs may vary. Thermokarst involves surface collapse due to the loss of contact between individual soil grains following ice melt. As a result thermokarst is constrained to ground containing at least excess ice. Field data indeed show that the plains highly affected by thermokarst in the Mackenzie delta, N.W.T. typically contain excess ice (Mackay 1966). In contrast, the minimum ice content in RTSs depends on the minimum amount of water (following the thawing of the ground ice) required to saturate the soil and produce flows that will continuously remove material from the scarp face and further down the floor of the slump in order to avoid the piling up that material in front of the ice face. This amount of ice is thus dependent upon the properties of the material and particularly, its liquid limit. In most cases however, massive ice have been reported in retrogressive thaw slumps (fig. 7), indicating that the amount of supernatant water is well above the liquid limit.

Figure 7. Massive ice in a retrogressive thaw slump. Debris flows can be observed on the slump floor. Image credit: Brian Moorman.

3.5.3. Material type and liquid limit

To understand the behavior of a material upon thawing, some parameters must be taken into consideration such as material type and liquid limit. As such the ability of a thawed block of ice-rich soil to flow over a distance depends on the liquid limit of the soil itself. It is known that in the case of retrogressive thaw slumps in the Mackenzie delta (Mackay 1966) and Herschel Island (Bouchard 1974), silt and clay sized materials and massive segregated ice are the dominant constituents within the soil profile although sand and gravel sized material also occur as bands within the ice in some instances such as in Thetis Bay (de Krom 1990). In the area of the Sand Hills moraine on southwest Banks Island, Lewkowicz (1987) noted that the retrogressive thaw slumps occurred in silty sand with occasional boulders. Silty clay underlying a layer of sand was also observed in an ablating head scarp downstream of Fort Simpson, N.W.T., by Pufahl and Morgenstern (1980). In Mayo, Yukon territory, Burn (1982) reported several metres of fine-grained sediments (2% sand, 75% silt, 23% clay) overlying coarser materials at depths > 12m. The liquid limit, also known as the upper plastic limit, refers to the water content at which a soil changes from the liquid state to the plastic state. In other words, the liquid limits marks the threshold at which a material will start to flow as a fluid instead of a plastic like in many glaciers. However, because a change in state usually occurs progressively in a soil, the boundary limit is often arbitrary. In a study of retrogressive thaw slumps on Herschel Island, de Krom (1990) notes that the liquid limit for the sediments that make up the studied slumps (mainly clay and silt with occasional thin layers of sand, peat and/or till) range between 30 – 40 %. However, because of the extensive presence of massive ice in the study area, the ice content of the material exceeded the liquid limits measured for the sediments.

3.5.4. Size of RTS

In the low arctic such as the Mackenzie delta area, RTS are typically 50 – 100 m in diameter with a scarp face between 2 and 7 m high. However, many RTS are polycyclic and extend over several hectares (French 1996). A polycyclic nature implies the formation

28 of a new RTS within the floor of an older one (active or stable), with the new one often establishing a similar morphology to its parent (fig. 8). As such, several generations of RTS within the same slump is common (Mackay 1966; Wolfe et al. 2001). Polycyclicity usually reflects the incomplete degradation of massive ground ice within the slump headwall (Mackay 1966). In the high Arctic, RTS tend to be broad and shallow. They rarely exceed 2 m in height but often possess a diameter of over 200 m (French 1974b; Lamothe and St-Onge 1961). Such disparities reflect the annual rate, depth, and duration of thaw between low and high Arctic regions (French 1996).

Figure 8. Polycyclic retrogressive thaw slump. The outline of the former scarp face can be seen in the middle of the photo. The older slump floor is now covered with vegetation. A newer RTS is developing inside the former RTS (bottom of picture). Image credit: Brian Moorman.

3.5.5. Retreat rates of RTS

Retrogressive thaw slumps are generally short-lived but rapidly developing landforms with the greater part stabilizing 30 – 50 summers following their inception. Thaw slump retreat rate varies and is highly dependent upon ground ice content. In southern Alaska, Wallace (1948) reported bank retreat rates of 0.06 – 0.19 m yr-1. In the Mackenzie delta region, annual retreat rates are inferred to range between 2 – 5 m (Mackay 1966) although higher rates have been noted in areas where wave action contributes to the erosion of the slump base (Lewkowicz 1985). In other areas, headwalls 10 – 20 m high have retreated at rates up to 9.6 m/yr and 30 m in exceptional circumstances (de Krom 1990; Lantuit et al. 2005). The sizes of RTS vary considerably from region to region. In the silty morainal terrain of eastern Banks Island, RTS display a regional density of 0.5/km (French and Egginton 1973) and typically exhibit retreat rates similar to those of the Mackenzie delta although a maximum of 15.5 m/year was documented for one slump in 1983-84 (French 1996). However, RTS in this area appear to remain active for a much shorter period (~ 12 – 15 years) and slumps have affected more than 15% of the coastal terrain. On a thaw slide in eastern Melville Island, NWT, Heginbottom (1978) reported retreat rates of 0.52 m (100h)-1. Herschel Island displays many of the largest as well as the highest frequency of RTS in the Canadian Arctic (de Krom 1990). Retreat rates on the island range between 0.3 and 2 m/yr and estimates of sediment volumes eroded due to RTS have shown to vary substantially with 360 000 m3 for the 1970-2004 period alone (Lantuit and Pollard 2005). In a study of thermokarst development in the Yukon Territory, Burn (1982) stated that the retreat rate of an ablating face relies upon the ground ice content and the amount of energy incorporated at the thawing surface from the atmosphere, given by McRoberts and Morgenstern (1974) as:

F V = (2a) + f TCL

Where V is the retreat rate, F is the energy flux at the surface which is largely determined by the amount of incident solar radiation, T is the difference in temperature of the ground from 0°C, Cf is the heat capacity, is the ice content and L is the volumetric latent heat of the soil. According to McRoberts and Morgenstern (1974), CfT is negligible compared to L, therefore:

F V = (2b) L

However, based on field measurements near Mayo, Yukon Territory, Burn (1982) found that maximum retreat rates occurred on days that were overcast and had rain. This latter event was likely due to the saturation of the upper layers of the ground profile due to heavy rainfall, followed by a reduction in shear strength of the headwall material.

3.5.6. Summary of RTS on Earth

On Earth, retrogressive thaw slumps are dynamic features of the periglacial environment and point to a disturbance in the thermal regime of ice-rich permafrost. As discussed above, the latter can be due to various factors resulting in the exposure of massive ground ice to atmospheric conditions and solar radiation. Such factors generally include active layer detachments and/or the slumping or erosion of the layer of material protecting the underlying ice. The consequences include the ablation of the ice and the continuous retreat of the ice face until stabilisation is reached. This stabilisation can be caused by the depletion of the ice or the complete coverage of the ice face due to slumping. In the case where undercutting of the scarp occurs due to stream or wave erosion, a reactivation of the RTS can occur, leading to the formation of a polycyclic slump. While rain and sublimation can contribute to the growth of a RTS, it is primarily the interaction between the atmosphere coupled with solar radiation and the ice face that will dictate the growth of a RTS.

CHAPTER 4

4. Mars

In his writings about a century ago, Percival Lowell popularized the possibility of water flowing in canal systems on the surface of Mars. This perspective changed in the 1960’s when the Mariner spacecraft found no evidence of canals. What it discovered instead, and probably of a much greater importance, are surface temperatures at which liquid water cannot remain stable (Ingersoll 1970). Inferences that Mars and Earth had similar climatic conditions early in their geological history (Haberle 1998) and the idea that Mars once had a warmer and wetter climate are not new (Carr 1996; Kargel 2004). Not surprisingly, early imagery of Mars revealed the presence of landforms that appeared to have formed under the work of running water (Pollack et al. 1987). But more importantly, the discovery of landforms similar to ground ice features on Earth gave an early indication that Mars’ climate has, at least at some point during its geological history, been able to sustain environmental conditions required by freeze and thaw processes. These landforms include possible pingos (fig. 9), polygonal patterned ground (fig. 10), thermokarst (fig. 10), lobate debris aprons and terrain softening (Squyres and Carr 1986). The search for more evidence of ground ice remains much of a priority today because of its significant implications for potential biotic activity (past or present) and a source of water. On the assumption that our knowledge of terrestrial ground ice processes and landforms is well founded, an understanding of the occurrence of near-surface ground ice on Mars, past or present, is essential in order to determine whether ground ice related landforms should exist or not on Mars (Chamberlain 2006). Like on Earth, the distribution of ground ice on Mars is dependent upon the current and past climate, availability of moisture, and the soil’s physical and thermal properties. As such, an understanding of each of these parameters is essential in the study of ground ice landforms on the Martian surface. This chapter looks at each of these components and reports on information available for each of them.

Figure 9. North is up. A). MOC R0100745. Degraded pingos in Athabasca Valles. Image size is ~ 1km across. B) MOC- E0500113. Possible pingos in a crater basin in Utopia Planitia. Image size is ~ 2 km across. Image credit: NASA/JPL/MSSS.

Figure 10. Thermokarst landscape in UP. North is up. (A) MOC-R0301203. Polygon pits and polygons. (B) MOC-E0200880. Polygon pits and polygons. (C) MOC-R0301203. Polygon pits, isolated depressions and polygons. (D) MOC-M1501497. Isolated and coalesced depressions. (E) MOC-E0200880. Polygon pits, depressions and polygons. (F) HiRISE TRA-000856-2256. Terraces inside alas-like depression. (A-E) Image credit: NASA/JPL/MSSS. (F) Image credit: NASA.

4.1. The Martian atmosphere and climate

Of all the planets that make up our solar system, Mars has the climate that is the most similar to that of Earth. Yet, Mars’ atmosphere is much thinner than Earth's with an atmospheric pressure of only 6 - 9 millibars compared to 1013 millibars for Earth. The atmospheric composition of Mars differs quite substantially from that of Earth’s with 95% carbon dioxide, 2.7% nitrogen, 1.6% argon, ~ 0.7% oxygen and other trace gases. With an atmospheric pressure as low as on Mars, liquid water cannot exist on the surface and would change directly from its solid form to its gas form, a process commonly know as sublimation.

Based on simple physical parameters, it is possible to determine the nature of the ice at the Martian poles. According to Lambert and Chamberlain (1978), with an atmospheric pressure of 6 millibars and assuming a ground temperature of < -123°C, CO2 ice can exist on the surface. A thin layer of frost covers both polar caps of Mars, estimated to be ~ 1 m thick (Sagan et al. 1972; Sagan et al. 1973). As mentioned above, the southern polar cap, slightly offset relative to the axis of rotation of Mars, exhibits extremely low temperatures

(-125°C in winter) and is therefore mainly composed of CO2 ice, and to a lesser extent,

H2O ice and clathrate at depth. In contrast, the boreal polar cap, larger than its southern counterpart, contains little to no CO2 ice because temperatures do not reach the freezing point of CO2 in summer (Costard 1993). Any CO2 that forms during the boreal winter would therefore disappear when warmer temperatures are reached. Depending on the amount of condensed carbon dioxide at the southern polar ice- cap, predominantly composed of carbon dioxide ice (dry ice), the atmospheric pressure may vary during the year (Carr 1996). Like on Earth, Mars’ polar ice caps extend and retreat according to the seasons. The northern polar ice cap (fig. 11) may reach latitudes as low as 65° during the northern winter. However, the longer and more intense southern winter results in a larger ice cap at the southern pole. This can be explained by the ellipticity of Mars’ orbit, which results in about 40% less heat received at aphelion than perihelion (Head et al. 2003). As a result, southern winters are longer and colder than northern ones. However, Mars’ precession is based on a 51,000-year cycle, which implies that in about 30,000 years time, southern winters will be shorter and warmer than they are now. Mars is at ~ 226,869,000 kilometres from the Sun (1.52 AU) compared to Earth (149,000,000 km). As such, Mars only receives ~ 44% of the sunlight received on Earth. Temperature gradients on Mars are very steep and the desiccated and overall granulous soil of Mars retains very little heat. This produces nighttime temperatures as low as -113°C at the equator whereas summer daylight temperatures can reach as high as 20°C or more, although the average surface temperature on Mars is about -55°C (Forget et al. 2003).

Figure 11: MOC2-270. Wide angle of the Martian north polar cap on March 13, 1999 (left) during early northern summer and January 26, 2001 (right). The north polar cap is ~ 1100 km in diameter. Image credit: Malin Space Science Systems

Direct measurements made by the Landers Viking 1 in Chryse Planitia (~22° N) and Viking 2 in Utopia Planitia (48° N) indicated air temperatures of -85°C - - 24°C and -81°C - -30°C respectively. Generally, the lowest temperatures are found at the South Pole during the winter where they can reach below -125°C (the freezing point of carbon dioxide). The Austral hemisphere also witnesses the highest temperatures at around mid-day during the summer where surface temperatures can reach 23°C but can drop down to about -100°C in winter (Forget et al. 2003). Early in the Martian morning and evening, ice crystal clouds form at high altitudes in the atmosphere, particularly near the summit of high volcanoes, such as Olympus and Pavonis Mons, in the Tharsis region (Kieffer et al. 1973; Wood 2001). In contrast, areas of low elevation witness early morning freezing fog.

Because Mars experiences seasons, the sublimation of CO2 and H2O ice at the polar caps into the atmosphere creates differences in air pressure. Although this change in pressure remains well below that required for water to exist at the surface, it is significant enough to induce movement of air from high to low pressure cells, which in turn generate global wind patterns on Mars (Fenton 2003). Analogous to the three Hadley cells on Earth, air rises and is transported in large circular motions on Mars. However, models suggest that the meridional topographic asymmetry of Mars results in one dominant cross-equatorial

(solstitial) Hadley cell except during a short transition period close to the equinoxes (Haberle et al. 1982; Haberle et al. 1993; Wilson and Hamilton 1996; Forget et al. 1999). As such, warm air rises at latitudes at which solar heating is the greatest, moves polewards, cools, and sinks. This circulation causes upwelling in the summer tropics, resulting in adiabatic cooling, and downwelling in the winter tropics, which causes adiabatic warming. As a result, the dominant Hadley cell transports heat from the summer hemisphere into the winter hemisphere and is associated with a significantly weaker cell at work in the summer hemisphere only (Richardson and Wilson 2002a). Due to the lack of a real stratosphere, which prevents vertical circulation of air, the Hadley cell on Mars can develop to a much higher altitude than those found on Earth. This global circulation of air is greatly altered by the rotation of Mars, which tends to displace the atmosphere with it. Similar to Earth’s Westerlies, air moving from north to south on Mars is entrained in an easterly direction, creating dominant west-east winds (Forget et al. 2003). In the upper part of the atmosphere where the effect of Mars rotation lessens, winds tend to be slower than the speed at which the planet is spinning therefore, creating a retrograde motion (east-west). Major wind circulation patterns on Mars are comparable to those found on Earth. For example, at the same latitude as Europe, winds on Mars blow mainly from the west but oscillate as low and high-pressure cells pass during the winter and autumn seasons. At the “tropical” latitudes of Mars, winds tend to blow in the same fashion as trade winds on Earth. However, diurnal oscillation of wind patterns is much greater on Mars because of the significant difference between day and night temperatures (Forget et al. 2003).

4.2. The water cycle of Mars

The Mars Atmospheric Water Detector (MAWD) on board the Viking orbiters was the first instrument to provide detailed observations of the water cycle on Mars (i.e. the distribution and transportation of water between the atmosphere and other reservoirs throughout the seasons) (Jakosky and Farmer 1982). The MAWD was followed by the Thermal Emission Spectrometer (TES) onboard the Mars Global Surveyor (MGS). The interesting observation, however, is that both datasets returned similar general patterns of water cycle. The highest water column abundances are located in the North Polar Region

37 during the northern summer with a relatively constant amount of water in the Martian atmosphere ranging from 10 – 20 pr-μm (precipitable microns of water). More recently, several models have tried to explain the Martian water cycle through the identification of the most important physical processes that determine water content in the atmosphere. Since the early days of atmospheric modeling, it was established that current atmospheric water content of Mars was principally controlled by the exchange of water between the atmosphere and the polar caps. A model from Davies (1981) and Jakosky (1983) showed that water from the poles sublimes, mixes diffusively across the latitudes and condenses on the surface at lower latitudes. Although Jakosky’s (1983) model implied water absorption by the regolith, resulting in a water reservoir over time, his conclusions supported the idea that polar processes control the atmospheric water content of Mars as well as other non-polar water reservoirs. Interestingly, more recent 3-D general circulation models (GCMs) have sustained this idea (e.g. Richardson and Wilson 2002b). As shown in figure 12 below (GRS map), an asymmetry in water distribution seems to exist on Mars with a larger concentration of water ice at the North Pole relative to the South Pole. As summarized by Chamberlain (2006), a number of reasons have been put forward to explain this north-south asymmetry. Early on, Davies (1981) argued that atmospheric water mixing during the southern summer due to large diffusivity values would be increased by dust storm activities associated with that season. Following the same idea, James (1985) stated that during the southern summer and spring seasons when the flow of global CO2 condensation is greatest, water transport to the north would be enhanced. It was explained that topography would play a role in water transport to the north as the latter is covered by seasonal CO2 frost for a longer period than the south (Jakosky 1983). The southern polar cap is topographically higher than the northern polar cap and has a lower CO2 frost point temperature, resulting in the retardation of seasonal frost formation which in turn results in a longer period of sublimation for any water ice present. North-south topographical asymmetries have also been inferred to enhance water transport to the north due to atmospheric circulation by creating a dominant southern summer Hadley circulation, the latter being independent of perihelion timing. Thus, the Hadley cell that is responsible for trade winds extensively controls water transport between

38 both hemispheres, which in turn influences climate, and eventually results in the increased formation of polar layered deposits in the north (Richardson and Wilson 2002a). As discussed above, only one Hadley cell exists on Mars and warmer global temperatures at perihelion increase water transport to the north as water is transported from the summer hemisphere to the winter hemisphere by the upper branch of the Hadley cell. At aphelion, water condenses out lower in the atmosphere due to cooler temperatures and thus reduces the amount of moisture that is transferred into this upper branch which in turn affects the ability of water to be transported to the southern polar cap at present (Clancy et al. 1996). For past water cycles, Jakosky and Carr (1985) proposed that during periods of high obliquity, ice would be deposited at low-latitudes. They estimated that the sublimation of polar ice would result in atmospheric saturation, precipitation and subsequent deposition of ice on the surface. More recent GCMs are in agreement with their results and indicate that the water cycle would result in the formation of ice on the surface at low latitudes during periods of high obliquity (Mischna et al. 2003; Mischna and Richardson 2005). By calculating the atmospheric water contents for past epochs, Mischna and Richardson (2005) found that at high obliquities, the atmospheric water content was not controlled anymore by the polar caps but rather by the new surface ice deposited at lower latitudes.

4.3. Models of ground ice in the Martian crust

Speculations about the presence of ground ice in the Martian crust started decades ago and were based on modeled temperatures and atmospheric water content (Leighton and Murray 1966). Interestingly, several lines of evidence acquired during the past few decades have supported such inferences. Evidence is provided for example by the gamma ray and neutron spectrometers, onboard the Mars Odyssey, which indicate high concentrations of hydrogen in the near-surface (upper metre) of Mars, representative of water ice (Boynton et al. 2002, Feldman et al. 2002; Feldman et al. 2004b) (fig. 12). This water ice, however, appears concentrated to high latitudes and in some areas to the upper parts of the mid- latitudes whereas equatorial regions appear to contain little to no water in the upper metre of the ground.

Figure 12. Water equivalent content on Mars based on hydrogen gamma-ray flux. Areas of ground ice appear as blue to purple. A trend that is in general agreement with Mars GCMs and ground ice models is that the equatorial regions should be desiccated at least to a few metres deep. The circle and square represent the landing sites of the Mars Exploration Rovers. The triangle indicates the location of Athabasca Valles. Image source: Boynton et al. (2002).

Other indications of ground ice in the Martian crust include high concentrations of craters surrounded by fluidized ejecta. Two types of impact crater ejecta exist on Mars. They are commonly known as unfluidized or fluidized. Unfluidized ejecta resemble typical lunar ejecta and display a non-cohesive character including many secondary craters, which implies emplacement from airborne objects ejected during the impact. In contrast, fluidized ejecta (also referred to as rampart craters) imply cohesive flow and tend to exhibit a morphology that resembles terrestrial debris flows or landslides (Costard and Kargel 1995). It has long been hypothesized that the fluidization of ejecta is a consequence of the melting or vaporization of ground ice and ground water caused by the impact (e.g. Carr and

Schaber 1977; Mouginis-Mark 1979; Wohletz and Sheridan 1983; Battisni 1984; Horner and Greeley 1987; Costard 1989a; Barlow and Bradley 1990; Baker et al. 1993; Barlow and Perez 2003). Although Schultz and Gault (1979) proposed the entrainment of Martian atmospheric gases, geologic evidence does not support atmospheric entrainment as the principal fluidization mechanism and does not explain variations in ejecta morphologies with latitude and crater size (e.g. Costard 1989a; Barlow and Bradley 1990; Baker et al. 1993). Rampart craters have thus been used as an indicator of ground ice in the Martian crust. The morphology of rampart craters varies significantly and probably point to varying volatile contents, physical state and/or vertical distribution of that volatile within the soil profile. Using Viking imagery, Costard (1989a) and Costard and Kargel (1995) collected data for 2600 rampart craters. Their results revealed a hemispheric asymmetry in ice content (or depth to ground ice table) between the northern lowlands and the southern highlands. Within this global variation, regional variations also became apparent in the distribution of rampart craters. Both Utopia and Acidalia Planitiae, located at the mouth of extensive outflow channels, exhibited very high concentrations of fluidized ejecta. They also noted that in young (post-heavy bombardment) terrains, variations in the percentage of rampart craters could be explained by the inability of smaller bodies to excavate beneath a surficial desiccated zone and into an ice-rich zone. By modeling the vertical variation in ground ice content based on rampart crater data, Costard and Kargel (1995) inferred that at latitudes < 40°N, the upper crust may be desiccated to depths of several metres to several hundred metres, again consistent with neutron spectrometry data. Figure 13 shows a three-layer model for Acidalia Planitia with a thin desiccated surface layer up to 60 m thick underlain by a ~900-1000 m thick ice-rich permafrost (mainly fluvial sediments), below which is found a less ice-rich (probably non-fluvial) layer. For latitudes > 40°, the authors note the same three-layer model applies since summer soil temperatures beneath the depth of diurnal thermal wave (10’s of centimetres to a metre) are below the frost point. As such, ground ice is normally stable at these latitudes below that point while the upper layer may undergo daily or annual cycles of ice enrichment or depletion. The

41 only difference between these two latitudinal bands is that the thickness of the upper desiccated zone may be less in the higher latitudes (Fanale et al. 1986).

Figure 13. Proposed ground ice model for Acidalia Planitia based on a study of rampart craters. The model is consistent with theoretical estimations of ground ice stability in the mid-latitude regions. (Costard and Kargel 1995).

Further geomorphological evidence pointing to near-surface ground ice in the Martian crust includes the presence of permafrost and ground ice related landforms such as pingos, thermal contraction polygons and thermokarst pits. Pingos are among the most visually intriguing landforms in permafrost environments and are defined as ice-cored mounds ranging between a few metres to more than 50 m in height and a few metres to hundreds of metres in diameter. The most common type of pingo forms following thermokarst lake drainage, permafrost aggradation and the freezing of water expulsed by hydrostatic pressure. In a study of Utopia Planitia, Mars, Soare et al. (2005) identified landforms similar to terrestrial pingos and inferred that these pingo-like mounds possess an ice-core that is still extant under ~ 1 m of soil. Similar landforms were identified by Burr et al. (2005) and Page and Murray (2006) in the equatorial region of Athabasca Valles (fig. 9) and the Cerberus plains although most of the latter ones are in a degradational state. However this would be consistent with thermal models with predict that ground ice would tend to decay at low latitudes (Kreslavsky and Head 2000; Mustard et al. 2001) and should therefore be non-existent in the upper metres of the ground at low latitudes including the equatorial regions.

Polygonal patterned ground is also common in the northern regions of Mars, particularly in Utopia Planitia. Polygons in permafrost are known to result from the thermal contraction of the ice-bearing ground due to sharp drop in temperatures. Several workers suggest that the polygons in Utopia Planitia are thermal contraction cracks underlain by ice wedges (e.g. Mellon 1997; Seibert and Kargel 2001; Mangold 2005). Thermokarst have also been identified in mid-latitude areas of Mars such as Chryse Planitia and the Circum-Chryse outflow channels, Utopia Planitia, the volcanic regions surrounding Olympus Mons, the southern highlands near Argyre Planitia and the cratered highland regions of Deuteronilus and Protonilus Mensae. (e.g. Sharp 1973b; Anderson et al. 1973; Gatto and Anderson 1975; Theilig and Greeley 1979; Costard 1987; Kargel and Strom 1992; Costard and Kargel 1995; Costard and Baker 2001; Seibert and Kargel 2001; Chapman et al. 2003; Mangold 2003). A recent analysis of high resolution Mars Orbiter Camera (MOC) and High Resolution Imaging Science Experiment (HiRISE) images suggests that thermokarst landforms in Utopia Planitia are recent (late Amazonian) and that excess or massive ground ice may still persist in areas unaffected by thermokarst degradation (Wan Bun Tseung et al., 2006). Besides geomorphological evidence, several thermal models have also been used to predict the extent of ground ice in the Martian crust (Mellon and Jakosky 1993; Mellon and Jakosky 1995; Mellon et al. 2004; Schorghofer and Aharonson 2005). One fundamental assumption of thermal models is that the atmosphere can diminish or replenish water in the subsurface through diffusive contact although other processes may form ice at or below the surface through ice migration within the soil profile or through atmospheric precipitation. In two different works, Mellon and Jakosky (1993, 1995) used thermal models to determine the distribution of ground ice on Mars under, 1. current conditions (Mellon and Jakosky 1993) and, 2. past obliquities (Mellon and Jakosky 1995). Under past varying obliquities, the 1995 one-layer model (in which it is assumed that the thermal properties of the ground were uniform with depth) suggested that planet- wide, ground ice would be stable at obliquities as low as 32°. For the present epoch, the model indicated that ground ice would be stable only at latitudes > 40° at a depth of about 100 metres and shallower polewards. Similar results were obtained by Mellon and Jakosky

(1995) who suggested that permanent ground ice is expected at latitudes > 40°. In such regions, ground ice is expected to occur below about a metre or even a few tens of centimetres of dry soil. Mellon and Phillips (2001) also reported that due to atmospheric condensation the Martian permafrost consisting of porous soils in contact with the atmosphere would be ice-rich at mid- and high latitudes. Following these results, Mellon et al. (2004) used a two-layer model (where the thermal properties of the soils were changeable) to determine ground ice distribution at present. Their findings showed that the high thermal inertia properties within the ice-rich layer allowed the ice to be stable much closer to the surface (within 1 metre), as previously demonstrated by and Farmer and Doms (1979) and Paige (1992). However, like in previous studies, the occurrence of ground ice close to the surface resided primarily within the mid- to high-latitudes (fig. 14).

Figure 14. Color scale indicates depth to ground ice. Clear regions indicate that ground ice is unstable in the upper metres of the ground. (Mellon et al., 2004)

Similar results were obtained by Schorghofer and Aharonson (2005) who also used a two-layer thermal model that involves water vapor diffusion to the atmosphere. These conclusions imply an atmosphere with a water content of 10 – 20 precipitable microns (pr-

m) which primarily results from the seasonal exchange of water with the residual polar caps and the latitudinal transport of that water from polar to non-polar regions (Haberle and Jakosky 1991). Other obliquity models also indicate that during periods of high obliquity (> 30°) when insolation is much more intense at the poles, water ice is stable near the surface down to the lower mid-latitudes (Mischna et al. 2003). Increased solar radiation at the poles releases great amounts of water vapor, thus increasing the amount of atmospheric water vapor content and, consequently, the humidity (Jakosky and Carr 1985). As a result, during these periods, the latitude of subsurface ice stability moves closer to the equator, and within the soil profile, closer to the surface (Mischna et al. 2003). Furthermore, Mars climate models indicate that ice is removed from the poles and deposited on the surface at lower latitudes where it may last throughout the year (Richardson and Wilson 2002b; Haberle et al. 2003; Mischna et al. 2003; Richardson et al. 2003). As a result, each high-obliquity period would remove tens of metres of ice from the poles, transport and deposit all this material, comprised of water ice and dust, in the mid- latitudes. Due to the higher year and summer average polar temperatures and denser atmosphere during high obliquity, water movement in the atmosphere is predicted to be rapid (Mischna et al. 2003). With the decrease of obliquity, the stability zones of surface ice shrink toward the poles and ice in the uppermost layer of the regolith tends to sublimate, thus, creating a buffering ice-poor crust for the in-depth ice. Furthermore, the decline of obliquity to < 21° results in significant drops of the atmospheric pressure as CO2 is locked in the both polar caps (Head et al. 2003). In summary, theoretical studies of ground ice depth are in general agreement and suggest that the equatorial regions of Mars are ice-poor, at least in the upper metres of the ground since ground ice is unstable at those latitudes. In contrast, mid- to high-latitudes should allow the presence of ground ice close to the surface.

4.4. Physical characteristics of the Martian fines

The Martian soil has always been thought to be composed mainly of small grains. Evidence for this was found a few decades ago through optical reflectance polarimetry with

45 telescopes (Dollfus 1957, 1958). Results at that time indicated small grains no larger than 50 m. However, recent analyses of the Martian surface based on optical polarimetric and photometric analysis were compared with the thermal inertia coefficients of Mars’ surface obtained by the Orbiter MARS-5 and the Viking orbiters. The results showed that the surface of Mars is covered by small grains ranging between 20 – 45 m except in very localized areas. Below this upper layer, radiometry indicated the presence of larger fragments of material such as rock slabs and stones (Dollfus and Deschamps 1986). The fine grains on Mars are inferred to be the result of aeolian processes primarily. Once sand- sized particles are available, the high-speed winds are able to produce saltation of the particles and cause impact fracturing, resulting in a wide range of grain sizes between 5 – 50 m (Smalley and Krinsley 1979). As documented by Dollfus and Deschamps (1986), subsequent in-situ analyses by the two Viking Lander cameras and soil samplers lead to the following conclusions: 1. Fine particles ranging from 10 - 100 m make up most of the Martian soils (fig. 15). 2. This property characterizes most of the Martian surface except where ice is present at the poles. 3. The smallest grains may cluster in clumps or stick to the surface of rocks and boulders through cohesion. 4. Unlike the Lunar surface, which tends to be rough, limited cohesion and atmospheric effects tend to produce smoother surfaces. 5. The grains are primarily composed of highly oxidized basaltic material to produce ferric oxides and hydrated clay minerals with water ice (McCord et al. 1982). 6. Aeolian processes are responsible for producing fine-grains by impact fracturing.

Recent results derived from thermal inertia observations in Melas Chasma indicated particle sizes ranging between < 10 m to up to 900 m (Pelkey et al. 2003). This larger range was explained by the fact that smaller dust particles are easily blown by wind whereas larger particles are more difficult to entrain from the canyon system. Several characteristics of Martian fines have been proposed since the early days of space exploration and the above-mentioned characteristics represent the most commonly used in

46 the scientific literature. However, despite the agreement between data returned by the Mars Exploration Rovers (MERs) and previously inferred characteristics, it is important to note that these values should only serve as a general guideline since the characteristics of the Martian soils are still poorly constrained.

Figure 15. Image PIA08619. Image of the Opportunity’s tracks in the Martian dust. Note the fineness of the dust. Image credit: NASA/JPL-Caltech/Cornell.

4.5. Thermal properties of the Martian permafrost

The thermal properties of the Martian soil are largely derived from thermal inertia analyses. The thermal inertia coefficient I is a measurement of the product of the compaction with bulk conductivity and is typically defined as a combination of the bulk material properties of thermal conductivity K, density , and heat capacity C such that I = (KC)1/2 where I is mainly dependent on K since both and C can only vary by a small factor between reasonable geologic materials and can thus be considered constant (e.g. Neugebauer et al. 1971; Kieffer et al. 1972; Kieffer et al. 1973). In turn, K is controlled by

47 particle size, porosity and pore connectivity in the top few centimetres of the surface (Putzig et al. 2006). Low values of I typically represent rapid adjustments to radiation heating fluctuations such as in highly loose material made up of micron-sized particles. High I values represent slow adjustments to variations in radiation heating and are usually indicative of hard bedrock (e.g. Dollfus and Deschamps 1986; Christensen 1986; Jakosky and Christensen 1986). As such, a loosely unbonded material made up of micron sized particles will commonly produce low I values around 2 10-3 cal/cm2 sec1/2 K whereas hard bedrock will typically produce I values higher than 50 10-3 cal/cm2 sec1/2 K (Dollfus and Deschamps 1986). Table 2a and 2b provide global values for the thermal and physical properties of the Martian soils. The geothermal heat flux (q) of Mars has not been measured yet but theoretical values range between 20 – 45 mW/m2 while a value of 30 mW/m2 is typically used (Johnston et al. 1974; Fanale 1976; Toksoz and Hsui 1978; Toksoz et al. 1978; Davies and Arvidson 1981; Stevenson et al. 1983; Franck and Orgzall 1987; Schubert and Spohn 1990; Mellon and Phillips 2001; Heldmann and Mellon 2004). Because thermal conductivity is affected by atmospheric pressure, the lower Martian pressures play a role in determining the conductivity of a soil. Although calculations of conductivity for unconsolidated soils on Mars do not exist, Mellon and Phillips (2001) assumed a 25% increase in the thermal conductivity for every 1 MPa increase based on calculations for the thermal conductivity of sandstone. Similarly, density also affects thermal conductivity. Presley and Christensen (1997b) calculated that at Mars pressures, thermal conductivity only increases slightly with density and can be expressed by k() = 0.01 + 6.4 x 10-6 W m-1 K-1 for densities of 700 – 1600 kg m-3. Following these results, Mellon and Phillips (2004) calculated only a 50% increase in thermal conductivity between the surface and 300 m depth. As such, variations in density should not affect thermal conductivity significantly when considering the upper few metres of the ground.

Table 2a Icy soil Icy soil Dry soil Property Parameter e e Units (CO2 ice) (H2O ice) (fine dust) Soil density 2274a 2018c 1650c Kg/m3 Specific heat c - 1287c 837c J/(kg K) capacity Thermal K 0.5b 2.4 0.045 (I2/c) W/(m K) conductivity Gravity g 3.71 3.71 3.71 m/s2 Geothermal flux q 30 30 30 m W/m2 Latent heat L - - 1.98 x 105 d J kg-1 Thermal diffusivity k - - 0.184 x 10-6 d m2 s-1

Table 2b Property Parameter Value Units J m-2 s-1/2 Thermal inertiaa I 250 K-1 Albedoa A 0.25 Density 1650 Kg/m3 b Porosity 0 0.4 (40%)

Table 2a (modified from Heldmann and Mellon 2004). a Weast (1986). b Kravchenko and Krupskii (1986). c Mellon and Phillips (2001). d Aguirre-Puente et al. (1994). Table 2b (modified from Mellon and Phillips, 2001). a Average global values for Mars based on Mellon et al. (2000). b Chosen to be consistent with surface soil densities estimated from radar and Viking Lander data (Moore et al. 1987).

CHAPTER 5

5. Results and discussion

Modeling ice-melt on Mars, particularly in the subsurface, remains challenging since most of the required parameters such as the soil thermal and physical properties, the geothermal flux and the orbital variations of Mars are either unknown or poorly constrained. Ground ice models for Mars also remain approximate and the proportions of ground ice as well as the depth to the ice table may vary largely depending of the geologic, atmospheric, climatic conditions and latitudinal position. Contrasting views among workers only add to the complicated task of modeling the thermophysical processes acting on ground ice on Mars. On Earth, studies are often accompanied by precise data recorded or obtained in the field, thereby enabling one to produce accurate and precise results, often for specific locations. On Mars, however, such results are not viable due to technological constraints and results therefore remain general and opened to further improvement. Nevertheless, such studies have proven fundamental in the understanding of Mars and form the basis of more detailed and accurate studies of the Martian geomorphology. Estimates of surface and subsurface heat flows on Mars exist but consider the entire thickness of the Martian lithosphere and mantle and are based on parameters such as strength of the lithosphere, radioactive elements and magnetization of the crust (e.g. Arkani-Hamed 2002; Ruiz et al. 2006). While these results may be significant in understanding the early thermal and magnetic state of the Martian crust and lithosphere on a global scale, they remain insignificant when considering processes (freeze and thaw) occurring primarily within the upper few meters of the ground. As such, determining the thermal characteristics of the shallow subsurface remains essential to the understanding of permafrost processes. Using existing estimated and calculated parameters for the Martian soil, this study identifies several factors that may have contributed to the formation of retrogressive thaw slumps on Mars. The first part (section 5.1.) of this discussion focuses primarily on heat conduction in the ground and argues about the possibility of RTS formation under current conditions on Mars. More specifically, a general argument about the potential for ground

50 ice degradation is made, followed by a more detailed analysis of temperature fluctuations at depth. Other processes are also discussed and evaluated as potential formation mechanisms for retrogressive thaw slumps on Mars. The second part (section 5.2.) elaborates on heat conduction results in the ground for past climates when Mars was warmer and wetter than today and draws on modeled and estimated temperatures for the last high obliquity period. Studies indicate that past climatic conditions may have been more favorable to freeze and thaw conditions under which the formation of RTS may have occurred. Such conditions include higher atmospheric pressures and warmer temperatures produced during high obliquity. Section 5.3. discusses the uncertainties associated with the results obtained in section 5.2. and section 5.4. elaborates on a sensitivity test performed in order to account for variances in the controlling parameter for thaw penetration in the ground. The last two sections (5.5. and 5.6.) of this chapter elaborate upon geomorphic evidence that suggest the presence of RTS on Mars.

5.1. RTS formation under current climate

Because Mars’ atmosphere is dry, it is reasonable to argue that near-surface ground ice may decay over time through diffusive loss to the atmosphere and may eventually result in the formation of a backwearing or downwearing ground ice landform such as a retrogressive thaw slump in the first case or thermokarst in the second. Indeed, theoretical and experimental estimates of vapour diffusivities in the Martian regolith indicate that under current conditions, ground ice exchanges vapor with the atmosphere (Flasar and Goody, 1976; Mellon and Jakosky 1993). This view is supported by the correlation between the geographic boundaries of subsurface ice (based on ground ice models) and ground ice stability (GRS data) for the present climate (Boynton et al. 2004; Mellon et al. 2004). However, as calculated by Mellon et al. (2004), under current Martian conditions and a dry atmosphere, a steady-state sublimation and diffusive loss of water vapor to the atmosphere (assuming 40% porosity) would result in the removal of about 400 μm of ice at 1 m deep (assuming a surface temperature of -53°C) in a little more than 3000 years. This suggests that seasonal fluctuations in the depth of the ice table are almost imperceptible. It

51 is also important to note that these values represent upper limits and are likely to be less due to the buffering of diffusion by adsorbed water and seasonal fluctuations in atmospheric moisture (Mellon et al. 2004). However, even considering a high diffusive loss of water vapor through the regolith and into the atmosphere, the formation of RTSs normally requires a thermal disequilibrium in the thermal regime of the ground. As such, the ability of a thaw front to penetrate to the ice table is essential and this latter requirement is largely dependent upon temperature fluxes at and below the surface.

5.1.1. Surface and subsurface temperatures

Spanovich et al. (2006) document surface and near-surface (1.1 m from the ground) temperature measurements taken during the hottest period of the southern hemisphere summer (known to exhibit the highest surface temperatures due to it close orbital position to the sun at perihelion) by the Opportunity and Spirit Mars Exploration Rovers (MERs) at the equatorial regions of Meridiani Planum (2°S, 6°W) and Gusev Crater respectively (15°S, 185°W) (fig. 16). Data returns from the first 65 sols (Martian days) show that in both locations, above zero daily surface temperatures occurred between ~10h00 and 16h00 LTST (Local True Solar Time) at Meridiani Planum and between ~11h30 and 15h00 LTST at Gusev crater. Peak surface temperatures reached 21.85°C at about 13h00 LTST at Opportunity’s landing site whereas the highest surface temperature recorded at Spirit’s landing site reached 6.85°C at about 14h00 LTST. Night surface temperatures dropped to a minimum of -98°C at Meridiani Planum and -90°C in Gusev crater although it is thought that lower temperatures were witnessed by the MERs but not recorded during the night (Spanovich et al. 2006).

Figure. 16. Landing sites of Opportunity (circle) and Spirit (square). Image credit: Ralph Aeschliman.

In contrast, near-surface atmospheric temperatures in Meridiani Planum reached a peak of -8.15°C at with a temperature profile that almost mirrors the surface temperature profile where near-surface temperatures rise slowly shortly after surface temperatures but fall at the same rate as surface temperatures. It is at about 16h30 LTST that surface temperatures fall below atmospheric temperatures and remain so until the following morning. In Gusev crater, a maximum near-surface atmospheric temperature of -13.85°C was reached shortly after landing. Like in Meridiani Planum, near-surface temperature profiles mirror surface temperature profiles except the highest temperatures are witnessed about one hour later in Gusev Crater. Figures 17a and 17b show that below zero temperatures are reached quickly after 16h00 LTST in Meridiani Planum and after about 15h00 in Gusev crater (Spanovich et al. 2006). Similar temperatures were also measured by the previous Viking (Tillman 1988) and

Pathfinder (Schofield et al. 1997) missions although the latter two primarily returned air temperature data.

Figure 17a. Surface and near-surface (1.1m above ground) temperatures taken by Opportunity. The temperature measurements are scattered over a period of 65 sols. Source: Spanovich et al. (2006).

Figure 17b. Surface and near-surface (1.1m above ground) temperatures taken by Spirit . The temperature measurements are scattered over a period of 65 sols. Source: Spanovich et al. (2006).

Because the Martian atmosphere is essentially transparent to solar radiation (i.e. atmospheric gases are ineffective at absorbing incoming shortwave radiation), it is the ground temperature that controls surface and near-surface temperatures through the release of sensible heat and conductive heat transfer (i.e. atmospheric gases are relatively good absorbers of long wave radiation emitted by the surface). If the surface temperature reached above zero temperatures for about 4 - 5 hours with a peak of 21.85°C (in the case of Meridiani Planum near the equator), it is then reasonable to assume that the ground temperature a few centimetres below the surface was above 0°C for at least a few hours. On Earth ground temperatures are usually 2 - 4°C higher than air temperatures close to the surface. It is known that in the subsurface, the amplitude of diurnal temperature variations decreases exponentially and is given by:

= ()I /c P / (3)

where P is the oscillation period, I the thermal inertia, the density and c the heat capacity. Based on thermal properties of the Martian soils, the diurnal skin depth is in the order of ~ 1 cm (Schorghofer et al. 2002) to ~ 6 cm with a common value of ~ 2.5 cm for a typical Martian regolith ( = 1650 kg/m3 and c = 837 J K-1 Kg-1) (Mellon and Phillips 2001). As a result, only the first ~ 2 cm of the ground can reach above zero (or melting) temperatures (assuming that the ground temperatures are close enough to 0°C so that only a few hours of above zero temperatures are sufficient to induce a raise in the ground temperature to the melting point of water). However, these first few centimetres of the ground is likely to be desiccated unless this dry layer has been removed by erosional processes. In order to account for more accurate thermal states at depth, temperature values at any depth z can be determined based on heat conduction properties for a steady-state condition given by (Williams and Smith 1989):

Tz = Ts + ()QG /K z (4)

where Ts is the surface temperature and QG is the amount of heat that flows by conduction and is given by:

QG = -K (dT/dz) (5)

where K is the thermal conductivity and (dT/dz) is the geothermal gradient (Williams and Smith 1989). While subsurface temperatures can thus be determined, it is important to note that in reality the heat conduction process is more complex because variations in the soil thermal regime (including the active layer) always occur due to changing natural conditions. Considering a sinusoidal surface temperatures variation and using known/estimated surface temperatures on Mars, the thermal properties of the Martian soil and a heat conduction model documented by Williams and Smith (1989), it is then possible to determine the

56 general features of the subsurface thermal regime. These features belong to that layer that witnesses annual variations (such as the depth of the seasonal wave). Because variations with depth in the thermal and physical properties of the Martian regolith are poorly constrained, it is assumed here that both parameters are uniform with depth. Ingersoll, Zobel and Ingersoll (1954, pp. 45-57) thus indicate that the temperature at any depth z is given by:

1/2 z() /2k 1/2 Tz(),t = T z + Ase sin t z (6) 2k

where = 2/P, P is the period of the wave (one year) in seconds, As is the amplitude of the surface temperature wave, t is the time in seconds, k is the thermal diffusivity, z is the depth in metres, and Tz is the mean surface temperature (Williams and Smith 1989). It must be noted that in the case of Mars, P is longer than 365 days since Mars is further from the sun and results in a longer revolution time. Furthermore, a Martian day (sol) is slightly longer than on Earth and lasts 24h39min35sec (Earth time). As a result, timekeeping on Mars can be problematical and a typical “Martian calendar” does not exist. However, the values given in this section are the most commonly used. As such, a Martian year lasts 668 sols or 687 Earth days. In order to avoid complications, the time scale for Mars used in the heat conduction model comprised a Martian year of 668 sols with each sol lasting 24 areohours (which takes into account the extra 39min35sec compared to Earth). The thermal diffusivity value used in the heat conduction equation is 0.184 x 10-6 m2 s-1 and is based on table 2a. The calculated thermal diffusivity of a typical Martian regolith is similar that of dry clay soil (40% porosity) on Earth. Surface temperatures were obtained from measurements taken by the MERs and cover the hottest period of the Martian year as well as the regions that would witness the highest temperatures on Mars. As mentioned in section 4.3, near-surface ground ice in the equatorial regions of Mars is inferred to be unstable and should only be found at several metres deep. However, to take into account the possibility of ground ice closer to the surface, as indicated by

57 potential ground ice landforms (pingos) in the equatorial regions of Athabasca Valles and the Cerberus Plains (Burr et al. 2005; Paige and Murray 2006), this study considers an upper limit for ground ice depth in the equatorial regions and assumes a depth of 2 -3 metres to the top of the ice-rich layer.

5.1.2. Heat conduction results

The following graphs show various features of the ground thermal regime at both Meridiani Planum and Gusev Crater, including surface and subsurface daily and hourly temperatures for both locations. The surface temperatures required in the heat conduction equation were obtained from data returned by the Mars Exploration Rovers. Because timekeeping for the MER missions does not follow the Martian calendar, the corresponding daily average surface temperatures (required in equation 6) for each calculated daily temperature cycle had to be determined. This was achieved by calibrating the recorded surface temperatures published with the raw data from the MERs. In the case of Opportunity in Meridiani Planum, the sol during which the highest surface temperature was recorded, sol 1 of the MER mission, corresponds to sol 167 in figure 18 (daily surface and subsurface temperatures over 1 Martian year). As for Spirit, which landed in Gusev Crater, the sol during which the hottest surface temperature was recorded, sol 1 of mission, corresponds to sol 167. Note that the results presented below are for the southern hemisphere summer, known to be the hottest due to the position of Mars at perihelion. Section 5.1.2.1. discusses results obtained based on published data in the scientific literature whereas the results in section 5.1.2.2. were calculated using the raw data returned by the MERs and obtained through the Planetary Data System. Differences between the two datasets include:

1. Disparities in daily maximum surface temperatures recorded. 2. A more extensive coverage (1 Martian year) of surface temperatures when using the PDS data compared to 180 sols of coverage in the published literature.

3. A calculated annual surface temperature (using PDS data) for section 5.1.2.2. instead of a published (Athena Science Payload team) annual surface temperature. 4. Differences in the annual and daily temperature amplitudes again involving calculated values based on PDS data instead of published values.

5.1.2.1. Results based on published data

In this section, Tz in equation 8 is taken as -53°C. This value was obtained from the Athena Science Payload (on board the MERs) team and is consistent with values calculated by Forget et al. (2003). Figure 18 indicates that at Meridiani Planum at the equator, average daily surface temperatures can reach above zero temperatures as high as about 8°C for 95 sols before rapidly falling down to below freezing temperatures. Similarly, above zero temperatures can penetrate the subsurface up to 22 cm and such average temperature can persist for 30 days before falling again below zero degrees Celsius. While such temperatures lasting a few days may cause the thawing of ice-rich permafrost, it is known that the upper metre of the ground in this region is desiccated and should therefore witness no thaw or melting. These rapid falls in temperature can be attributed to a relatively low thermal diffusivity of the ground, which in turn is representative of a low heat capacity. Unlike Earth where vegetative and snow cover have a profound effect on the energy exchange between the subsurface and the atmosphere, on Mars the lack of such coverage combined with the lack of a thick atmosphere and clouds contribute to the loss of heat during nighttime.

Figure 18. Meridiani Planum (2°S, 6°W). Daily surface and subsurface temperatures over one Martian year (668 sols). Average surface temperature is taken as -53°C.

The same non-thawing trend can be observed in figure 19a which indicates that throughout most of the Martian year, average temperatures remain well below 0°C except for a thin upper layer (0.22 m) that witnesses thaw. Another interesting feature of the temperature regime shown in figure 19a is the quick stabilization of temperature variations at about 14 m deep. This depth is likely to represent the annual thermal skin depth, or, the depth zero annual amplitude (the depth at which temperatures variations changes over a year do not exceed 0.1°C). The depth of zero daily amplitude for sol 1 of mission in Meridiani is at 0.4 m.

Figure 19a. Annual temperature envelope for Meridiani. Blue lines represent temperature variations over 1 Martian year. Red lines represent temperature variations over one sol. Upper right rectangle is shown in next figure.

Figure 19b. Zoom-in of upper right corner of figure 19a. Black horizontal bars represent temperature range during the hottest sol of the MER mission. Diagonal bar (left center) represents maximum temperatures over one Martian year.

A closer look at the hourly temperature variations over the hottest sol of the year (fig. 19b) indicates that the surface temperatures exhibit pronounced variations, reaching a high of 22°C. However, a quick dampening of that variation can also be observed and temperature hourly temperature fluctuations rapidly decrease at 0.3 m deep. This thin upper layer can be characterized the depth of the active layer in this part of Mars. Subsurface temperature profiles were also plotted for the sol (137) during which the minimum surface temperatures were recorded and published. Figure 20 shows these temperature variations and indicates that, unlike those calculated for sol 1 of mission in, subsurface temperatures fluctuated with an amplitude of almost 125°C. However, only surface temperatures are able to reach above 0°C and any geomorphological phenomena therefore remains in the cryospheric zone and under frozen conditions. The depth of zero daily amplitude for sol 137 of mission can be observed at about 0.5 m. Temperatures at

62 greater depths were also plotted (fig. 21) and indicate that they quickly fall to less than - 50°C at 5 m deep. Such low temperature can be attributed to the fact that sol 137 of mission corresponds to the beginning of the southern fall season on Mars.

Figure 20. Meridiani Planum (2°S, 6°W). Hourly surface and subsurface temperatures for sol 137. Average surface temperature is taken as -53°C. The graph shows surface and subsurface temperatures for the sol during which the minimum surface temperature was recorded (sol 137 if mission) by the MER. Dashed line represents the depth (0.05m) of calculated values closest to the 0°C.

Figure 21. Surface and subsurface temperature profile for Meridiani on sol 137 of mission showing average temperatures.

Surface and subsurface temperature variations were also derived for Gusev Crater at the landing site of Spirit. In contrast with Meridiani Planum where daily average surface temperatures may reach up to 8°C, Gusev Crater exhibits daily average surface temperatures that remain below 0°C throughout the Martian year as shown in figure 22a with an depth of zero annual amplitude of 13 m.

Figure 22a. Annual temperature envelope for Gusev Crater. Blue lines represent temperature variations over one Martian year. Red lines represent temperature variations over one sol. Upper right rectangle is shown in next figure.

Figure 22b. Zoom-in of upper right corner of figure 22a. Black horizontal bars represent temperature variations over the hottest sol of the MER mission. Diagonal bar (left center) represents maximum temperatures over one Martian year.

The same cold trend can be observed in figure 22b suggesting that any geomorphological activity remains primarily in the frozen zone. This interpretation is supported by the calculation of hourly surface and subsurface temperature fluctuations for the hottest sol (sol 1 of mission) at Gusev Crater (fig. 23). On sol 1 of mission (peak of southern summer), only surface temperatures can reach up to 7°C. Even then, this temperature can only be maintained for a few hours and above zero temperatures can propagate only to 5 cm deep as indicated in figure 23 and are therefore restricted to the desiccated zone. In contrast, sol 180 of mission (beginning of southern fall) (fig. 24) shows surface and subsurface temperatures that are well below zero degrees Celsius with a low of about -90°C. The depth of zero daily amplitude for sols 1 and 180 are at 0.4 and 0.45 m respectively.

Figure 23. Gusev Crater (15°S, 185°W). Hourly surface and subsurface temperatures for sol 1 of mission. Average surface temperature is taken as -53°C. Dashed line represents the depth of calculated values below which temperature remain below 0°C.

Figure 24. Gusev Crater (15°S, 185°W). Hourly surface and subsurface temperatures for sol 180 of mission. Average surface temperature is taken as -53°C.

5.1.2.2. Results based on raw MER data

In this section, results derived using the heat conduction equation (6) used an average annual surface temperature of -34.8°C and -37.81°C for Meridiani Planum and Gusev Crater respectively, compared to an average annual surface temperature of -53°C for both locations in the previous section. A larger temperature range was observed using the raw MER data primarily because the data encompasses temperature recordings of one complete Martian year in contrast to 180 sols in the published literature. More specifically, it is Gusev Crater that exhibits the largest temperature variation although not significantly different from that of Meridiani Planum. As such, the maximum and minimum surface temperatures recorded in Meridiani Planum were 26.77°C and -96.43°C respectively and 30.07°C and -105.03°C. Like on Earth, Gusev Crater’s larger temperature amplitude could be due to its more southerly position compared to Meridiani, which is close to the equator and therefore witnesses less variations in temperature throughout the year. The temperature profile for Meridiani (fig. 25) shows that surface temperatures can reach up to a maximum of 27°C over the year and that above zero average daily temperatures can penetrate down to 1 m deep (also representative of the depth of the active layer) and remain at 0.9°C for about 15 sols before falling well below the freezing point and reaching the depth of zero annual amplitude at about 13 m deep. Likewise, in Gusev Crater, average temperatures at 1 m deep can reach up to about 1.3°C for several days (fig. 26) but quickly fall below zero degrees Celsius below that depth.

Figure 25. Annual temperature envelope for Meridiani Planum (2°S, 6°W). Blue line represents temperature variations over one Martian year and red lines represent temperature variations over the hottest sol of the year.

Figure 26. Annual temperature envelope for Gusev Crater. Blue line represents temperature variations over one Martian year and red lines represent temperature variations over the hottest sol of the year.

5.1.2.3. Thaw of ice-rich permafrost

Results obtained in sections 5.1.2.1. and 5.1.2.2. are in general agreement with ground ice stability models for Mars and indicate that under current conditions, a thaw front could reach to a maximum of 1 – 1.1 m deep in the low latitudes. Consequently, this heat

70 flux is therefore restricted to the desiccated zone and the ice table should witness no thaw. Such a shallow penetration of the thaw front in this region of Mars is probably due to the low conductivity of the ground combined with the inability of the soils to retain heat. This in turn results from the short “warm” (> 0°C) hours witnessed at the equatorial regions, even during the summer season. In addition, freeze back following a thaw episode would be accelerated by both the occurrence of cold air temperatures and from the permafrost below. However, like Earth, Mars witnesses weather changes regularly, sometimes on a diurnal basis leading to variations in daily temperatures and intensity of insolation. On Earth, ground surface temperatures can sometimes reach up to about 20°C above the air temperature close to the surface, leading to snowmelt in winter although air temperatures remain below 0°C (Forget et al. 2003). For Mars, this phenomenon is shown in figures 17a and 17b where average surface and near-surface temperatures for consecutive days exhibit drastic differences. As a result, these changes in temperatures at the surface should be reflected in the subsurface, suggesting that if surface temperatures can reach up to 27°C (20°C above the maximum average surface temperature recorded during the MER mission), subsurface temperatures should therefore reach the melting point of ice. In order to account for this possibility, the surface temperatures used in the heat conduction equation were raised by 20°C. However, the results indicate that in the subsurface, temperatures remain well below 0°C with little to no change compared to the results obtained using the original values for surface temperatures. For example, on the hottest sols in Meridiani Planum and Gusev Crater, an increase in surface temperature of about 20°C only led to temperature variations within the upper 0.5 m of the ground while below this depth, temperatures quickly decreased to < -20°C. This phenomenon is likely due to the low thermal diffusivity of the Martian soils combined with the presence of extremely cold permafrost below the surface. As such, any increase in surface temperatures would be dampened at depth. The global asymmetry in ground ice distribution on Mars probably contributes in further decreasing the potential for ground ice thaw in these regions. This same boundary dichotomy (topographical asymmetry that separates old rough terrains and young, post- bombardment, flat terrains) is likely to be the primary factor that accounts for the non-

71 occurrence of periglacial processes in the southern low latitudes and is coincidental with the results presented in sections 5.1.2.2. and 5.1.2.2. Indeed, terrains south of the dichotomy boundary tend to exhibit a larger number of impact craters that possess unfluidized ejecta. In contrast, the northern plains of Mars, particularly Acidalia and Utopia Planitia possess numerous rampart craters indicative of ground ice at relatively shallow depths. While these rampart craters may account for ground ice emplacement in the late geological history of Mars and the potential for retrogressive thaw slumps to form under current conditions, a lack of detailed and precise ground temperature measurements does not allow an accurate prediction of RTS formation in these regions. If we assume for one moment, based on ground ice models for Mars, that the depth to the top of the ice table is at 1 m deep, the likelihood for a thaw front to penetrate to this depth remains unlikely in the mid- to high latitudes. This interpretation is supported by the heat conductions results obtained above which indicate that the thaw front can only penetrate to 1 metre deep in the equatorial regions and that this depth should be shallower as one moves away from the equator. Furthermore, the thermal regime of the upper layers is dependent upon exchanges of heat and moisture between the atmosphere and the ground surface as well as the thermal properties of the ground. In turn, the climate and the surface energy balance (thus the geographic location). However, the interaction between these two parameters is complex and the surface temperature regime is therefore affected by more than just the geographic location (Williams and Smith 1989; French 1996). On Earth, other influencing factors include local surface conditions such as type of vegetative cover and depth of snow cover, both of which having a profound effect of energy exchange between the ground and the atmosphere. It is these site-specific conditions that will produce a microclimate which in turn will control ground thermal conditions because differences in surface cover will provoke a variation on heat energy exchange between the ground and the atmosphere (Williams and Smith 1989). On Mars, however, these factors are undoubtedly absent and should therefore discard the “warming effect” that snow/vegetation would have on the ground during the winter. In the unlikely case of ice-melt due to a thaw front, several other factors should be taken into consideration.

First, the surface atmospheric pressure on Mars is typically 500 – 600 Pa and in both cases, below the triple point of water (610 Pa). This factor alone would restrict the existence of liquid H2O on the surface and would therefore prevent the flowage of saturated material as required in the formation of an RTS. Second, following the assumption that the surface atmospheric pressure may reach above 610 Pa for a minimum amount of time for liquid water to exist on the surface and eventually carry sediments away, it is important to take into account that thawing above freezing temperatures can only occur for a limited number of days (which in turn differs one season to another and from one year to another) during the Martian year and although above 0°C values may be reached, temperatures do fall well below the freezing point for many hours over one sol. As a result, meltwater (if capable of existing at atmospheric conditions) would refreeze very rapidly due to cold temperatures. Third, similar to Earth, it is well known that while surface and subsurface temperatures can reach above 0°C for a few hours, air temperatures can still remain below freezing on Mars (Costard et al. 2002). This view is supported by air temperature measurements made by the MERs (see section 5.1.1.) which indicate that peak air temperatures during the hottest days of the mission reached a maximum of about -8°C in Meridiani Planum and -13°C in Gusev Crater. In such a case, any water produced and exposed to the atmosphere through the melting of ground ice would quickly readjust to atmospheric conditions which, on Mars, would lead to both sublimation and/or very rapid freezing due to cold temperatures and a low atmospheric pressure. Fourth, because the Martian atmosphere is dry, the water content in the soil may diffuse out of the ground primarily because of the high ground-atmosphere water density gradient at the moment of thaw in the ground. This in turn would result in the removal of water from the ground before the latter reaches saturation. However, Martian soils are primarily composed of fine-grained sediments and this diffusion would be slow. Nevertheless these four factors, individually or in combination, are likely to restrict the formation of RTSs under current Martian conditions, at least under typical Earth processes. As such, the formation of RTS on Mars currently should be constrained to sublimation processes exclusively following an active layer detachment. As stated above,

73 under current conditions on Mars, water and ice cannot exist on the surface except at the polar caps and exposed ice would tend to sublimate at mid- and low-latitudes. Under such conditions and considering a dry atmosphere, the exposure of a massive ice face caused by the erosion of the overlying material (wind or slumping) (Mellon and Phillips 2001) would rapidly result in the retreat of the ice wall and further slumping of the overlying material. Figure 27a below illustrates a theoretical ice face exposed to Martian atmospheric conditions. Sublimation of the massive ice would result in headwall retreat and further slumping of the overlying material. Assuming a constant sublimation rate and using equation 6 it is then possible to calculate the retreat rate V of an ablating ice face on Mars using theoretical values of thermal properties of the Martian soil such that

F V = (7) L

is 0.0185 m/day or almost 2 cm per day. Similar to Earth, the stabilization of RTS on Mars is a function of the depletion of the ice or the balance between the amount of debris supply and the removal of that debris from the base of the ice wall. However, because liquid water is not produced during ground ice sublimation, the efficiency of flowage is greatly diminished. As a result, the continuous ablation of the ice face will result in the gradual accumulation of debris at the base of the ice wall and a slow blanketing of the ice face as shown in figure 27b, c, d. On Mars the amount of debris required to insulate ice from atmospheric conditions is dependent upon the diurnal and seasonal thermal skin depths or the depth of the active layer. As mentioned above, under present atmospheric conditions and a typical Martian regolith the diurnal skin depth is in the order of centimetres to a metre or more in the case of less dense material (Jakosky and Christensen 1986) although very low subsurface temperatures and the high conductivity of ice will “wick” heat away from the surface and further enhance the stability of ground ice by dampening the “warming” effect caused by surface conditions (Schorghofer and Aharonson 2005). Furthermore, Mars is known to

74 witness regular severe dust storms, sometimes global, capable of burying an ice exposure rapidly under several centimetres of dust.

Figures 27a, b, c, d. A hypothetical ice face exposed to currents condition on Mars. Sublimation to the atmosphere would lead to the gradual retreat of the ice face followed by gradual slumping of dry debris on the ice face, eventually ending in a complete coverage of the ice face.

Based on heat conduction results above, at least ~ 2 cm of soil is required to insulate ground ice from surface conditions. The amount of slumped material could thus be determined by knowing the size of the ice face (area) and the angle of repose of the material. This amount, however, will in turn be affected by the flowage capacity or the cohesiveness of the material. Sand will tend to produce a dry debris flow whereas cohesive material such as duricrust, which refers to a weakly cohesive soil material formed by chemical precipitation and hardening of sedimentary bodies, will tend to fall in blocky pieces. On Mars, estimated values for the angle of repose of a typical surface material (moderately dense dry soil composed mainly of fine sandy grains) varies between 27° and 39° with a commonly used value of 35° (e.g. Perron et al. 2003). As mentioned above, figure 27 illustrates a hypothetical ice face with dimensions similar to RTS headwalls on Earth. This conceptual model assumes an ice face with a length (A – B) of 10 m at and angle of 80°, an overburden of 1 m (Fig. 27a) and a constant ablation rate resulting in the slumping of the dry overburden. The model depicts the exposure of an ice face to Martian conditions. Figure 16a shows the ice face subject to solar radiation and sublimation. Figure 27b and 27c show partial covering of the ice face due to the slumping of the overburden following the retreat of the ice face. Figure 27d shows the

76 complete covering of the ice face when the latter reaches the angle of repose (35°). In the case of a complete stoppage of the growth of the RTS, subsequent slumping events may restart the evolution of the RTS if the ground ice has not been depleted completely during the previous cycle, resulting in a polycyclic RTS. It is worth noting, however, that this model illustrates ideal conditions, uniform thermal properties and assumes that no outside processes affect this process of ice wall retreat and slumping. In real conditions, micro and macro–scale factors such as differences in material type within the overburden or aeolian processes such as dust storms may alter the landform or affect the processes leading to its formation or degradation.

5.1.2.4. Melting of ground ice at temperatures below 0°C

As discussed above, the calculation of surface and subsurface temperatures for the equatorial regions indicates that under present conditions, thaw can only occur occasionally in the upper centimetres of the ground. Because this upper layer should be desiccated, the temperatures derived suggest no thaw of the deeper ice-rich permafrost. The formation of retrogressive thaw slumps, solely dependent upon the melting of excess ice, is therefore restricted to other components of ground thaw not involving above zero temperatures. Similar results are likely to be the same in higher latitudes because of the inability of a thaw front to reach the ice table due to colder temperatures as one moves polewards. However, it is essential to take into account other potential influencing factors that could account for the melting of near surface ground ice and the presence of liquid water in the near surface. The most probable scenario involves the presence of salts in the ground. This could shift the triple point of water to lower temperatures. Such a phenomenon is well known to permafrost scientists on Earth where saline water may keep the ground unfrozen even a few degrees Celsius below freezing. Inference about the presence of salts in the Martian permafrost is not recent and this is in great part due to the observation of duricrust at both Viking and Pathfinder landing sites (e.g. Mutch et al. 1977; Binder et al., 1977; Kargel 2004). Brass (1980) argued that if one or more salt species were present, the freezing point depression could be decreased to as low as -50°C. However, as reported by Mellon and Phillips (2001), the problem is the very high amount of salts required: 15% – 40% by

77 weight is required for melting to occur (with an ice table temperature of about -70°C). In contrast, seawater on Earth contains ~ 3.5% salt by weight and freezes at ~ -1.85°C (Horne 1969). Although Kuzmin and Zabalueva (1998) inferred that brine pockets could exist in the Martian regolith, there have been no documentation of such high concentrations yet. Furthermore, the 8% - 15% sulfates and 0.5% to 1.5% chlorides and trace bromides calculated for the Viking landing sites would only decrease the freezing depression point by a few degrees (Kuzmin and Zabalueva 1998). While salt deposits could potentially result in the formation of near-surface water by lowering the freezing point, a more substantial account and observation of such deposits is required.

5.1.3. Alternative RTS formation mechanisms under current conditions

Igneous melting on Mars is considered as one of the ways by which liquid water could form in the subsurface. This process is considered by some as how catastrophic outflow channels (e.g. Ogawa 2003; Kargel 2004) and gullies may have formed (e.g. Malin and Edgett 1999; Gaidos 2001; Heldmann and Mellon 2004). Intrusive volcanism has been shown to produce heat flows able to melt ground ice and form subsurface water reservoirs (e.g. McKenzie and Nimmo 1999; Ogawa et al. 2003). Calculations made by Gaidos (2001) indicate that a magma intrusion 1 km below the base of an ice-bearing ground could bring the melting depth of ground ice at about 200 m from the surface. However, at such depth and without a continuous intrusion of geothermal heat, the water would refreeze. As calculated by Gulick (1998), the time required for a volcanic intrusion to melt a 2 km thick permafrost filled with ice (40%) is on the order of several thousand years or less assuming the presence of a 50 km3 intrusion and a temperature of about -35°C for the permafrost. This would have resulted in a subsurface reservoir just above the volcanic intrusion and below an ice-rich permafrost still thermally controlled by atmospheric and solar conditions (Fanale et al. 1992). Additionally, modeling of hydrothermal convection within Mars’ mantle indicate that subsurface water can be brought up close to the surface but only to a few hundred metres from the latter. This value, while still significant in terms of groundwater formation, remains too low to have implications in the near-surface (top few metres) melting of ice.

Other thawing mechanisms such as aeolian deflation could also be invoked but involve mechanisms not well understood. Alternative scenarios could imply a more important role of sublimation than occurs in the terrestrial analogue landscapes or perhaps salt hydrate deposition and dehydration, but these processes also would require a higher humidity and/or an abundance of liquid not unlike Earth. In summary, the alternatives involve a hypothetical landscape evolution with no known complete terrestrial analogue and little to non-existing documentation.

5.2. RTS formation under past (warmer) climate

While it is known that today Mars is a dry and cold planet with surface ice locked only at the polar caps, many geomorphic indicators such as gullies, channels, erosional valleys, salt-rich evaporite beds and other features point to a warmer and wetter environment than today. Such a warm and wet environment would have been able to sustain liquid water in large quantities, possibly enough to support an ocean in the northern lowlands (Head et al. 1999; Malin and Edgett 1999). It is now widely accepted that such a warm and wet Mars could have been possible at periods of climate warming driven by orbital variations (e.g. Touma and Wisdon 1993; Costard et al. 2002; Kargel 2004).

5.2.1. Orbital variations and climate change

As shown by various stratigraphic records, large-scale changes in Earth’s climate are driven by small intermittent variations in the Earth’s eccentricity, obliquity and precession over timescales of 10,000 – 100,000 years (Milankovitch 1941; Hinnov 2000). Likewise, the modelling of the Martian obliquity indicates that its orbit and axial behaviour has varied with timescales comparable to Earth’s, although these variations have been more pronounced on Mars, partly due to the stabilizing effect of the moon on the Earth (Chamberlain 2006). While it is inferred that Earth’s obliquity has oscillated between 22° and 24.5° over the past 10 million years, Mars’ obliquity has ranged between 14° to 48° (Head et al. 2003), sometimes reaching an extreme of ~ 60°, also referred to as chaotic obliquity (Jakosky et al. 1995). In comparison, the eccentricities of both planets have

79 witnessed proportionally different behaviours. The eccentricity of Earth’s orbit has ranged from 0 to 0.06 whereas Mars’ orbit has varied from 0 to 0.12. Such orbital variations have presumably resulted in significant changes in insolation and seasonality at mid- to high-latitudes on Mars (e.g. Ward 1992). However, it is now assumed that of the three orbital parameters, obliquity is the primary control on climate change. Changes in obliquity, which refers to the tilt of the axis of rotation with respect to the orbital plane, result in fluctuations in the latitudinal distribution of solar energy. Studies suggest that before five million years ago, Mars’ obliquity consistently exceeded 45° with a mean obliquity of > 35°. During the past 4 million years, Mars has witnessed a mean obliquity of ~ 26° with a range between 15° and 35° (fig. 28) (Kasting, 1991; Head et al. 2003). This is turn has principally affected the seasonal cycles of carbon dioxide and water in the Martian atmosphere. According to Head et al. (2003), Mars may thus have witnessed the most pronounced climatic variations over the past 10 million years of any planet in the Solar System. Recently obtained high-resolution images of the Martian surface coupled with older imagery reveal the presence of landforms associated with ground ice such as creep in young terrains which suggest recent climate changes on Mars. Terrains exhibiting various degrees of dissection, interpreted as thermokarst (Sharp 1973b; Costard and Kargel 1995; Wan Bun Tseung and Soare 2006) also seem to indicate that cemented surficial material partially disaggregated due to ground ice melting, evaporation or sublimation. The amount of dissection has in some instances been linked to the amount of ice deposition from the atmosphere during periods of high obliquity followed by degradation of that ice at low obliquities (Mustard et al. 2001; Milliken and Mustard 2003).

Figure 28. Martian obliquity cycles for the past 5 million years (Ward, 1992).

Based on model calculations of the behavior of the CO2 and H2O cycles at high obliquity, Jakosky et al. (1995) inferred that chaotic obliquity (~ 60°) would have had significant implications for the seasonal water cycle. The latter could be extremely rapid and remove the total thickness of polar deposits in a period as short as 10,000 years.

Furthermore, the CO2 trapped in the polar caps could be released into the atmosphere and therefore result in periods of warmer, more temperate climate throughout geologic time accompanied with higher atmospheric pressures. Another factor that makes Mars more susceptible to climate change is the ability of surface temperatures on the planet to drop below the condensation point of CO2, the main constituent of the atmosphere (Chamberlain 2006). It is known that at the present obliquity,

~ 25% of the atmospheric CO2 condenses onto the winter pole as CO2 ice creating seasonal variations in the atmospheric pressure of Mars. It is inferred that higher obliquities would result in significant variations and extensive ice caps. In contrast, during periods of low obliquity, colder temperatures in Polar Regions may form expansive CO2 ice caps, lowering the atmospheric pressure. Eventually, such changes would be reflected in fluctuations in the water cycle and content of the atmosphere (Chamberlain 2006).

Estimates of erosion rates show that there has been a fast decrease in Noachian (oldest period of Mars’ geological history- 4.6 -3.5 billion years- characterized by rimless, flat-floored impact craters and valley networks) terrains to exceptionally slow rates during the Hesperian (second oldest period: 3.5 - ~ 3 billion years) and Amazonian (~ 3 billion years to present) (Forget et al., 2003). This fast decrease of erosional rates seems to be consistent with a major shift in climatic conditions during the Noachian and coincides with the end of the heavy bombardment period, and a cold, dry climate similar to that of today. Golombek and Bridges (2000) argue that the Noachian abrasion rates on Mars appear to be consistent with the work of running water and therefore, a more temperate climate.

5.2.2. Melting of ground ice at high obliquity

Ground ice on Mars should be present close to the surface at mid- and high- latitudes. Based on geomorphological evidence, this ice could be in the form of excess or massive icy beds extending to tens of metres into the subsurface (e.g. Farmer and Doms 1979; Fanale et al. 1986; Squyres and Carr 1986; Paige 1992; Costard and Kargel 1995; Mellon and Jakosky 1995; Mustard et al. 2001; Costard et al. 2002; Soare et al. 2005; Wan Bun Tseung et al. 2006). Mustard et al. (2001) state that alterations in insolation due to variations in obliquity primarily, but also eccentricity and time of perihelion, may cause the removal of ground ice in the upper (2 –4 m) regolith over timescales of 100,000 years, particularly at latitudes >30°. As predicted by Touma and Wisdom (1993) high obliquity increases the atmospheric pressure above 611.73Pa and causes average surface temperatures to exceed 0.01°C, both parameters defining the triple point of water. As discussed above, under present atmospheric conditions the diurnal skin depth is in the order ~ 2cm or more. With an increasing ice or water content within the soil profile, the amplitude of diurnal, seasonal and annual skin depths (this being enhanced during high obliquity) should increase due to the high thermal conductivity of the H20 constituents. While the seasonal skin depth in terrestrial ice-rich permafrost may reach > 15 m (Williams and Smith 1989), the higher obliquity of Mars associated with more intense and longer insolation periods could result in a deeper penetration of the energy flux juxtaposed on the surface and melt near-surface ground ice. The modeling of heat transfer into the ground

82 supports this interpretation. Using the heat conduction equation (6) and estimated surface temperatures during high obliquity (Costard et al. 2002), subsurface temperatures were derived for a period of high obliquity (45°) as it may have occurred at summer solstice (Ls 270°) 100,000 years ago. Figure 29 is a reconstruction of the temperature profile for Mars at latitude 50°S for that period. The average surface temperature used here is 11°C. The most interesting feature of the graph is that average daily temperatures during that period were able to remain above 0°C to several metres deep, reaching a high 32°C at the surface. Hourly temperature calculations (fig. 30) support this argument and indicate that even over one sol, average temperatures persist above 0°C. On the assumption ground ice is present at less than one metre deep at high latitudes, a thaw front can thus reach the ice-rich zone and eventually cause ice-melt. This melting effect is likely to be accelerated by the latent heat release from the melting ice. The same process can be observed on Earth. However, the thawing of ice-rich permafrost would only be one of the results of such a climatic change. As discussed above, significant changes to the atmosphere including a higher pressure, an increase in water content and humidity and higher temperatures would in turn allow the existence of any liquid water resulting from the melting of ground ice on the surface. As mentioned by Costard et al. (2002), it is the southern mid- to high-latitudes that would witness the highest temperatures due to an increase in obliquity coupled with the position of Mars at perihelion. The consequences of such processes could lead to various results including the formation of liquid water in the ground and the saturation of that ground (depending on material type, pore size and amount of water produced). If pore pressure is increased to a point higher than the shear strength, an active layer detachment may occur since the thaw front would have a downward component and thus permafrost could still be extant at the moment of slumping. This permafrost, in turn exposed to solar conditions would witness thaw and eventually degrade. The formation of an active layer detachment could then trigger the formation of a retrogressive thaw slump, which would be able to survive due to high temperatures, pressures and high atmospheric water content, especially at the peak of summer. Furthermore at that period of the year, a high precipitation rate due to a high atmospheric water content caused by a major depletion of the polar caps could also result in

83 high pore water pressures in the ground, thus enhancing the likelihood of an active layer detachment.

Figure 29. Reconstruction of the annual temperature envelope for Mars at latitude 50°S during high obliquity (45°) about 100,000 years ago.

Figure 30. Calculated daily surface and subsurface temperatures for Mars over 1 sol at Ls 270° (summer solstice) during high obliquity (45°).

Ground ice can also be exposed to atmospheric conditions and/or be melted through mechanical or thermal erosion. This can be achieved due to the atmospheric changes implied by high obliquity cycles where liquid water should be able to run freely on the surface. The triggering of a RTS through mechanical (e.g. lateral stream erosion) or thermal erosion (heat transferred by flowing water) is not exclusively a Martian process since it is known that on Earth many retrogressive thaw slumps form through these processes (fig. 31). While atmospheric and ground ice stability models for Mars indicate that the environmental changes induced by high obliquity can lead to the melting of near-surface ground ice, predictions remain general, primarily due to the lack of data for past atmospheric conditions.

Figure 31. Retrogressive thaw slump triggered by lateral stream erosion and slope detachment. Image credit: Brian Moorman.

5.3. Uncertainties

Surface temperatures used in equation 6 for sections 5.1.2.1. and 5.1.2.2. were obtained from data returned by the Mars Exploration Rovers. While surface temperatures are in general agreement in both instances, a slight offset between minimum and maximum surface temperatures can be observed. As indicated in table 3 below, in the published literature, maximum and minimum surface temperatures recorded are 21.85°C and - 98.15°C respectively for Meridiani Planum and 6.85°C and -90.15°C respectively for Gusev Crater. In contrast, surface temperatures obtained directly from the MER raw data are 26.77°C and -96.43°C for the maximum and minimum temperatures in Meridiani Planum and 30.07 and -105.03 for Gusev Crater.

Published data Maximum temperature (°C) Minimum temperature (°C)

Meridiani Planum 21.85 -98.15 Gusev Crater 6.85 -90.15 Raw data Meridiani Planum 26.77 -96.43 Gusev Crater 30.07 -105.03

Table 3. Minimum and maximum surface temperatures obtained from published data and raw data.

The slight offset between the minimum temperatures for both locations can be explained by the fact that the published data included only the first 180 sols of the MER mission whereas temperatures derived from the raw data covered one Martian year (668 sols). In both cases (published and raw data), maximum temperatures fell under the same period of the year, corresponding to the beginning of the MER mission and should therefore be more or less similar. However, maximum temperatures for Gusev crater that show the greatest difference with an offset of about 23°C. Like in all Mars studies, discrepancies exist among researchers and the methods used to obtain results, particularly when dealing with unknown surface and subsurface temperatures. This purpose of this study was to make use of the most accurate data available in the published literature to investigate the thermal characteristics of the Martian ground and to compare these results with calculated values obtained from the raw data. Here, differences in the methods used to derive temperatures undoubtedly led to disparities in the results obtained. In all cases, however, readjustments of the calculated temperatures to reflect the published temperatures would only result in a lowering of overall temperature results and confirm that any thaw would occur in the upper desiccated layer.

5.4. Sensitivity test

Of the various parameters in the heat conduction equation (6), thermal diffusivity k is the main parameter controlling the depth of penetration of a heat wave. In order to account for variations in thermal diffusivity in Martian soils, a sensitivity test was done to determine differences in subsurface temperature variations. The k value used in the heat conduction equation (i.e. 0.184 x 10-6 m2 s-1: dry clay soil) represents an end member between a range of diffusivity values of materials commonly encountered on Earth. As such a thermal diffusivity value of 0.9 x10-6 m2 s-1 (wet sandy loam) was used to estimate the potential subsurface temperatures variations on Mars. Although wet sandy loam is not expected on the surface and near surface of Mars, it is assumed to be a diffusivity test value since it represents the other end member of the thermal diffusivity range in Earth materials. The test indicates that an increase in the thermal diffusivity also increases the depth of penetration of a thermal wave as shown in figure 32 and figure 33 for both Meridiani Planum and Gusev Crater respectively. As a result, above freezing temperatures are able to persist down to 0.5 m (Fig. 32) in Meridiani Planum compared to 0.22 m using the estimated thermal diffusivity value. Similarly in Gusev Crater, despite a similar increase in depth of penetration of a thermal wave, the maximum average temperatures still remained below freezing throughout the year (Fig. 33). The diffusivity test value was also used in determining temperature changes over the hottest sol of the year. The results show that over one sol in Meridiani, temperatures vary with a range of 29.8°C at the surface but rapidly dampen with little to no variations at 0.5 m below the surface, at which depth temperatures fall below the freezing point. A similar behavior was noted for Gusev Crater except temperatures remain under 0°C below 0.2 m depth. Overall, similar results were obtained using the raw data and they indicate that a thaw front would remain restricted to the upper desiccated zone in the low latitudes. The results from the sensitivity tests are in agreement with predicted models of ground ice stability for Mars and indicate that a thaw front at the low latitudes would not be able to reach the top of the ice-rich layer, estimated to be a ~ 3 m depth.

Figure 32. Sensitivity test for the annual temperature envelope for Meridiani Planum. The thermal diffusivity value in this graph is taken as 0.9 x 10-6 m2 s-1. The blue line represents temperature variations over one Martian year and red lines represent temperature variations over the hottest sol of the year. Black rectangle is described in figure 34.

Figure 33. Sensitivity test for the annual temperature envelope for Gusev Crater. The thermal diffusivity value in this graph is taken as 0.9 x 10-6 m2 s-1. The blue line represents temperature variations over one Martian year and red lines represent temperature variations over the hottest sol of the year.

Figure 34. Zoom-in of upper right corner of figure 32. Black horizontal bars represent temperature variations during the hottest sol of the MER mission. Diagonal bar (left center) represents maximum temperatures over one Martian year.

Figure 35. Zoom-in of upper right corner of figure 33. Black horizontal bars represent temperature variations during the hottest sol of the MER mission. Diagonal bar (left center) represents maximum temperatures over one Martian year.

5.5. Retrogressive thaw slumps in Athabasca Valles

The results obtained above indicate that under steady state conditions, ice melt should not occur. However, steady state conditions are rare if not absent on Mars and factors such as topography, aeolian activity and fluvial activity contribute to producing dynamic landscapes where erosion (either by wind or water) may result in the removal of the desiccated upper layer in certain areas of Mars. Only in situ analysis would be able to confirm the presence of an active or relict RTS on Mars. However, some features on Mars suggest that RTS could be present today. Figures 36 and 37 show two areas in Athabasca Valles (see figure 12, page 39 for location), the youngest known outflow channel on Mars. Numerous bowl shaped features with steep headwalls and what appears to be slump floors are present along the main channel of Athabasca Valles. Initial measurements indicate that

92 size (width) of the bowl shaped features in figures 36 and 37 range between ~80 m to ~500 m, similar to RTS on Earth.

Figure 36. Subset of HiRISE image PSP -002793-1885 (8.2°N, 154.4°E). North is to the left. Illumination is from the bottom left. Bowl-like features in Athabasca Valles. Disturbed floor at in the lower part of the image represents the floor of an outflow channel. In set is described in figure 38. Image credit: NASA/JPL/University of Arizona.

Figure 37. Subset of HiRISE image PSP-002437-1875 (7.2°N, 151.9°E). North is to the right. Illumination is from the top right. Slump features on the western side of the Athabasca Valles area. Slump feature denoted by red arrow is ~ 91 m across. Image credit: NASA/JPL/University of Arizona.

5.5.1. Features of the RTS in Athabasca Valles

In addition to the morphological similarity between the two types of landforms, the presence of polygons (Fig. 38) further indicates that the bowl shaped landforms in Athabasca Valles may have a permafrost origin. On Earth, polygons are usually associated with ice-bearing ground and form due to repeated sharp temperatures drops below the freezing point. This leads to the thermal contraction of frozen ground and subsequent cracking, releasing tensile stress. Over time, a network of polygons forms (Lachenbruch 1962). Polygons in the Athabasca Valles area of Mars have been calculated to have an average maximum diameter of ~ 25 m and are thought to be the result of thermal contraction cracking of ice-bearing ground (Burr et al. 2005). Among the most interesting features of the Athabascan slumps are small blocks (a few metres in size) of material located on the scarp faces (Fig. 38). Based on HiRISE images, these slab-like features

94 appear to be fragments of material that have detached from the surface and fallen onto the slump wall. Furthermore, the crisp edges of the slumps suggest that detachment of surface material is recent and perhaps ongoing. At the base of the scarps, small scale sand dunes can be observed and point to active aeolian activity in the region. Such activity could have significant impact on the erosion and exposure of ground ice and/or the burial of that ice.

Figure 38. In set of Figure 36 showing the details of the slump features. The bowl like feature (middle left is ~ 86 m across). Blue arrows point to polygons. Red arrow shows small meter scale sand dunes. Rectangle show detached overburden on the scarp face. Small circular depressions also appear on the slump floors. Note that the headwall character is similar to that of the RTS on Earth but mud flows are not evident on the slump flow as are produced by liquid water flow in RTSs on Earth. Image credit: NASA/JPL/University of Arizona.

5.6. Ground ice and RTS formation in Athabasca Valles

Athabasca Valles is located in the low latitudes and global Gamma Ray Spectrometry (GRS) data indicates that the ground in these regions should be ice poor (Boynton et al. 2002, Feldman et al. 2004). However, as pointed by Burr et al. (2005), GRS

95 maps are composed of ~300 km pixels and the RTS-like features shown in figures 36, 37 and 38 would only represent a tiny portion of a pixel. Therefore, the presence of ground ice in localized pockets could be possible. While morphology points to a permafrost origin for these features on Mars, it is interesting to note that the geomorphological history of Athabasca Valles is also in favor of processes that may have triggered the formation of RTS in the region. This involves water flow within channels. Flooding in this region of Mars is thought to have occurred in the past, with the most recent event dating back to 2 – 8 Ma (Burr et al. 2002b). Studies by MacEwen et al. (2005) suggest that Athabasca Valles could be two orders of magnitude older. However, this would still remain within the late Amazonian epoch. With multiple floods having flowed down the Athabasca channels (Burr et al. 2002b; Keszthelyi et al. 2004a), it is very likely that under freezing conditions, supersaturated soils would have resulted in the formation of massive icy beds in the region. Furthermore, the presence of numerous rampart craters, indicative of ground ice (Costard and Kargel 1995), predating the Athabasca floods in the region suggest that ground ice was even present before the Athabasca channels were carved. This interpretation is based on the streamlined forms on the main channel of Athabasca Valles, as observed by Burr et al. (2002b). The presence of pingo-like features in Athabasca Valles further points to the presence of ground ice in the area (Burr et al. 2005) although other mechanisms such as Jökulhlaups (glacier outbursts) (Maizels 1997) may have resulted in the formation of the pingo-like features in Athabasca Valles. The striking resemblance between the bowl shaped features on Mars and RTS on Earth (Fig. 39) suggest that they may have the same origin. Two scenarios for triggering the formation of RTS are implied here. First, flooding in the area could have involved lateral stream erosion (mechanical erosion) of material or thermal erosion of ground ice along the flow channels and caused the exposure or degradation of the ice. Such processes are common in terrestrial permafrost environments. Second, wind erosion could expose ground ice through the removal of surface material partly eroded by precedent flooding episodes. In both cases, the presence of polygons could have contributed in accelerating the slumping of surface material by acting as planes of weakness in the ground. The exposed ground ice to atmospheric conditions would result in the formation of RTS features similar to those

96 illustrated in figures 27a, b, c, and d. The absence of flow features on the slump floors further suggests that melting is/was not involved in the formation of these features.

Figure 39. Top image of retrogressive thaw slumps at Peninsula Point, N.W.T., Canada showing multiple (polycyclic) slumps. Image credit: Google Earth. Bottom image is HiRISE PSP -002793-1885 (8.2°N, 154.4°E). Image credit: NASA/JPL/University of Arizona.

What is of further interest is the fact that if the RTS-like landforms in Athabasca Valles were triggered by the most recent flooding episode, this indicates that permafrost processes were active during the recent past and that ground ice in this region may still be extant. In figure 36, the slump floors immediately beneath the scarp faces appear to be smooth compared to the distal parts of the slump. This inconsistency in surface roughness is

97 attributed to two or more episodes of ground ice degradation. Such a process is common in RTS features and is referred to as polycyclicity where older RTS have not been able to melt the entire ice body or where environmental conditions caused the stoppage of RTS growth. Subsequent events restarting the growth of these features result in the formation of younger RTS within their parent features (Fig. 40). The presence of numerous small circular depressions ranging from less than a meter to a few meters on the slump floors further indicates that more than one episode of ground ice degradation may have occurred. These depressions could be result of localized sublimation of ground ice that remained following an initial RTS formation.

Figure 40. Polycyclic retrogressive thaw slumps in Herschel Island, Canada. Blue arrows show delineation between old and young slumps.

CHAPTER 6

6. Summary and conclusions

In recent years, the exploration of Mars has seen an unprecedented amount of data returned by various missions. While these data have undoubtedly contributed to a better understanding of Mars, much remains unknown about its subsurface. This is due primarily to technological constraints that the scientific and engineering communities are now facing. As a result, theoretical and conceptual models are still heavily used to infer the physical and thermal characteristics of the Martian permafrost. The aim of this study was to determine whether retrogressive thaw slumps can or could have formed on Mars and is principally based upon heat conduction into the ground since it is a disruption in the ground thermal regime that primarily triggers the formation of a RTS. Retrogressive thaw slumps are among the most widespread geomorphic features in permafrost regions of Earth and are indicative of ice-rich ground. On Mars, reports of permafrost related features have so far been absent of retrogressive thaw slumps, indicating that if ground ice is present close to the surface, thaw temperatures are not able to penetrate deep enough to melt the ground ice. The results obtained in this study from the heat conduction equation (6) are in general agreement with ground ice stability studies for Mars. They indicate that although temperatures > 0°C can be reached for a few hours over one sol under contemporary conditions in equatorial regions, these temperatures can only reach a few decimetres or a maximum of one metre into the ground. With no surface erosion, this upper layer will likely be desiccated and therefore no melting of ground ice should occur. This argument, however, does not account for the potential for rapid surface erosion of the sediments caused by fluvial or aeolian processes. A reconstruction of the thermal regime for a past climate indicates that atmospheric and climatic changes due to high obliquity could result in a deeper penetration of the thaw front at high-latitudes where ground ice is close to the surface. Under such conditions, the thaw of ice-rich permafrost would then be probable. This argument is again supported by the observation of numerous degradational ground ice features in the northern plains of

Mars such as Utopia Planitia (e.g. Costard and Kargel 1995; Soare et al. 2005; Wan Bun Tseung et al. 2006) and Ares Valles (Costard and Baker 2001). In order to account for variations in the thermal properties of the Martian soils and in particular thermal diffusivity (the controlling factor in determining heat propagation in the subsurface), a sensitivity test was completed using a thermal diffusivity value that would encompass the whole range of materials that could potentially be found on Mars. As expected, the results of the sensitivity test indicate that despite a higher thermal diffusivity in the low latitudes of Mars, above freezing temperatures are likely to remain within the upper meter. The presence of features similar to terrestrial ground ice landforms and the cryotic state of Mars suggest that permafrost processes may have been dominant during the geological history of the planet. While this inference remains difficult to confirm without precise data about the Martian surface and subsurface, several lines of evidence further enhances the idea that freeze and thaw processes once occurred in the Martian history. Several bowl-like features (slumps) are present along the main channel of Athabasca Valles, one of the several outflow channels on Mars. Morphologically these features are similar to RTS on Earth and the fresh morphology of the slumps suggests that they postdate the flooding episodes that carved Athabasca Valles. Additionally, the geomorphological history of Athabasca Valles is in favor of ground ice emplacement during the late Amazonian when outflows may have resulted in the saturation of the sediments with water. Subsequent flooding episodes over thousands of years may have caused mechanical or thermal erosion of surface material and resulted in the degradation of ground ice. If the bowl-like features in Athabasca Valles are RTS features, this implies that ground ice was present in the recent past and may still be extant and could have significant implications for future Mars exploration missions to the low latitudes. In summary, under current temperature cycles and considering the presence of a desiccated layer at the surface of Mars and a slow ground ice diffusion, modern RTS should not occur. However, RTS could have formed in the past under warmer conditions. In contrast, considering fast surface sediment removal due to fluvial or aeolian activity as it

100 may have been possible in Athabasca Valles, RTS formation could potentially occur under current conditions at localized sites of sediment removal.

6.1. Future research

The results obtained in this study do not exclusively pertain to the formation of retrogressive thaw slumps since the thaw of ground ice in general can result in a suite of geomorphological features such as thermokarst basins, collapsed pingos and degraded ice- wedge polygons. Because of that, the methods used in this thesis can easily be transferred to other ground ice thaw studies in order to provide supportive arguments. The documentation of the thermal characteristics of the ground at the landing site of Opportunity and Spirit could also serve as a base reference for further exploration of these sites or others not only within the same latitudinal band but also in other regions of Mars. However, because heat conduction studies require a certain number of known parameters, future research and studies should include a more detailed and accurate understanding of surface and subsurface characteristics of the Martian soils. These characteristics could include a greater number and more precise measurements of surface temperatures as well as true readings of subsurface temperatures. Undoubtedly, high-resolution coverage of Mars will eventually take place with time and new imagery will help in better investigating the Martian surface. However, future rover-based missions should implement findings based on satellite imagery in order to corroborate or confirm inferences about potential ground ice features on Mars. Furthermore, accurately determining thermal characteristics of the ground such as thermal conductivity, heat capacity and thermal diffusivity should also be part of future missions’ mandates in order to achieve a better characterization of the Martian subsurface. This is because while surface and atmospheric processes are important in understanding Mars, a profound knowledge of the Martian subsurface is crucial since it is there that any future exploration missions are likely to find resources such as water and potentially, past or present signs of biotic activity.

101

Evidence for the presence of icy sediments in localized areas of Mars exists in the form of ground ice features. However, the challenge remains to extrapolate this information to larger areas of Mars. As such, this indicates the importance of considering the temporal and spatial scale of the multiple processes that shape the Martian landscape and their resulting geomorphic features.

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