Periods of active permafrost layer formation during the geological history of Mars: Implications for circum-polar and mid-latitude surface processes

ARTICLE IN PRESS Planetary and Space Science 56 (2008) 289 302 www.elsevier.com/locate/pss Periods of active permafrost layer formation during the geological history of Mars: Implications for circum-polar and mid-latitude surface processes Mikhail A. Kreslavsky a,b,, James W. Head a, David R. Marchant c a Department of Geological Sciences, Brown University, Providence, RI 02912, USA b Earth and Planetary Sciences, University of California, Santa Cruz, Santa Cruz, CA 95065, USA c Department of Earth Sciences, Boston University, Boston, MA 02215, USA Received 18 July 2005; accepted 2 February 2006 Available online 29 August 2007 Abstract Permafrost is ground remaining frozen (temperatures are below the freezing point of water) for more than two consecutive years. An active layer in permafrost regions is defined as a near-surface layer that undergoes freeze thaw cycles due to day-average surface and soil temperatures oscillating about the freezing point of water. A dry active layer may occur in parched soils without free water or ice but significant geomorphic change through cryoturbation is not produced in these environments. A wet active layer is currently absent on Mars. We use recent calculations on the astronomical forcing of climate change to assess the conditions under which an extensive active layer could form on Mars during past climate history. Our examination of insolation patterns and surface topography predicts that an active layer should form on Mars in the geological past at high latitudes as well as on pole-facing slopes at mid-latitudes during repetitive periods of high obliquity. We examine global high-resolution MOLA topography and geological features on Mars and find that a distinctive latitudinal zonality of the occurrence of steep slopes and an asymmetry of steep slopes at mid-latitudes can be attributed to the effect of active layer processes. We conclude that the formation of an active layer during periods of enhanced obliquity throughout the most recent period of the history of Mars (the Amazonian) has led to significant degradation of impact craters, rapidly decreasing the steep slopes characterizing pristine landforms. Our analysis suggests that an active layer has not been present on Mars in the last 5 Ma, and that conditions favoring the formation of an active layer were reached in only about 20% of the obliquity excursions between 5 and 10 Ma ago. Conditions favoring an active layer are not predicted to be common in the next 10 Ma. The much higher obliquity excursions predicted for the earlier Amazonian appear to be responsible for the significant reduction in magnitude of crater interior slopes observed at higher latitudes on Mars. The observed slope asymmetry at mid-latitudes suggests direct insolation control, and hence low atmospheric pressure, during the high obliquity periods throughout the Amazonian. We formulate predictions on the nature and distribution of candidate active layer features that could be revealed by higher resolution imaging data. r 2007 Elsevier Ltd. All rights reserved. Keywords: Mars; Past climate; Permafrost; Cryoplanation 1. Introduction Permafrost occurs on the Earth in vast high-latitude regions and local high-elevation areas (e.g., Washburn, 1973; Pewe, 1991), where the year-average surface temperature is below the water freezing point (e.g., Williams Corresponding author. Department of Geological Sciences, Brown University, Providence, RI 02912, USA. E-mail address: kreslavsky@brown.edu (M.A. Kreslavsky). and Smith, 1989). Muller (1947) defined permafrost on Earth as a soil or rock layer in which temperatures are below the freezing point of water for a continuous period of two years. The layer of frozen ground can be up to hundreds of meters thick (e.g., Brown, 1970). The upper meter(s)-thick layer of the ground in permafrost regions can undergo a yearly freezing and thawing cycle; this is known as the active layer (Fig. 1) (Miller and Black, 2003). In the spring, a thawing wave propagates downwards into the permafrost and in the fall and early winter, a freezing 0032-0633/$ - see front matter r 2007 Elsevier Ltd. All rights reserved. doi:10.1016/j.pss.2006.02.010

290 ARTICLE IN PRESS M.A. Kreslavsky et al. / Planetary and Space Science 56 (2008) 289 302 Fig. 1. Diagrams illustrating different configurations of permafrost and active layers. (a) Permafrost: the situation in which the day-average temperature is below the freezing point of water all year around, and no active layer is present. (b) (c) Active layer: the situation in which there is seasonal melting of water ice. During the warm season (b), surface temperatures exceed the freezing point of water, and a thawing front propagates down through the active layer. Then, during the cold season (c), surface temperatures fall below the freezing point and a freezing front (the frost table) propagates downward from the surface and the active layer progressively freezes. A second freezing front slowly propagates upward from the permafrost table. Later in the cold season the entire thickness of the active layer becomes frozen. (d) Active layer with talik: under some conditions the freezing front (frost table) does not reach the permafrost table in the cold season, and a confined layer of unfrozen ground (known as talik) is present year round. This could be due to surface temperature trends over short (year to year) or longer (millennial) time periods. wave propagates down from the surface, sealing and progressively freezing the active layer. Occasionally, freezing during early fall also commences upward from the permafrost table (Mackay, 1981, 1983). The frost table delimits the seasonal base of the active layer (the limit of the freezing front as it descends downward in the winter) and this usually, but not always corresponds to the permafrost table (the top surface of the underlying perennially frozen ground) (Martini et al., 2001). In cases where it does not reach the permafrost table, soil with liquid water can exist year-round in the intervening space and is known as talik. This situation occurs when the yearaverage surface temperature is close to 0 1C, and the permafrost is metastable or marginally stable. The permafrost table forms an impermeable barrier at the base of the active layer. The thickness of an active layer depends primarily on the temperature of the atmosphere, and secondarily on substrate heat conduction. Typically, active layer thicknesses are greater away from the pole; they are thinnest under insulating snow and thicker in areas of higher heat conduction. In some permafrost areas, such as the high-elevation parts of the Antarctic Dry Valleys (Bockheim and Hall, 2002), the day-average surface temperature never exceeds the freezing point of water, and a traditional active layer does not form (e.g., Marchant and Head, 2004, 2005). Dry permafrost is the special case where neither free water nor ice is present, although some moisture in the form of interfacial water may occur. Dry permafrost is thaw stable and accompanying dry active layers, if present, lack cryoturbation; thus, neither dry permafrost nor dry active layers will be further considered in detail here. A water-containing active layer in permafrost regions causes a number of specific surface modification processes and produces a variety of distinctive geomorphic features. These include cryoturbation (frost heaving), patterned ground formation (including sorted and unsorted circles, nets, steps, stripes and polygons), solifluction and gelifluction (gelifluction is associated with frozen ground; lateral flow produces sheets, lobes, benches), cryoplanation (reduction of the topography of a terrain surface by ice and active layer processes), thermokarst (variety of surface depressions caused by melting of ground ice), etc. (e.g., Washburn, 1956, 1973; French, 1976; Williams and Smith, 1989; Krantz, 1990). The scale of these individual features ranges from sub-meter to many hundreds of meters. Some specific types of features form on permafrost surfaces without an active layer, for example, sublimation polygons observed in the Antarctic Dry Valleys (Marchant et al., 2002; Marchant and Head, 2005). Geomorphologic features characteristic of active layer processes are abundant on Mars (Fig. 2). High-resolution images taken by the MOC camera onboard Mars Global Surveyor (Malin and Edgett, 2001) have revealed a variety of features and patterns, including polygonal ground. Most of the polygonal ground occurs at high latitudes. A number of researchers have analyzed these features and pointed to permafrost-related and active layer-related origins (e.g., Carr and Schaber, 1977; Rossbacher and Judson, 1981; Lucchitta, 1981, 1983, 1985; Squyres and Carr, 1986; Squyres et al., 1992; Carr, 1996; Mellon, 1997; Seibert and Kargel, 2001; Masson et al., 2001; Kuzmin et al., 2002; Kuzmin, 2005; Leverington, 2003; Mangold et al., 2002; Mangold, 2005). Among the variety of polygon

ARTICLE IN PRESS M.A. Kreslavsky et al. / Planetary and Space Science 56 (2008) 289 302 291 Fig. 2. An example of martian polygons morphologically similar to terrestrial ice-wedge polygons suggestive of the active layer processes. Small-scale knobby pattern superposed over the large polygons is interpreted to be sublimation polygons formed in the absence of an active layer. This place is located on the floor of a crater at 621N221E. Portion of MOC NA image E21/01593, illumination is from lower left. morphologies on Mars, there are examples strikingly similar to terrestrial ice-wedge polygons (Fig. 2). Such polygons occupy the floors of many high-latitude impact craters. These morphological similarities suggest a similar origin that appears to require formation of an active layer. Similarly, Mangold (2005) noted slope morphologies strikingly similar to terrestrial solifluction lobes which also require seasonal thawing of an outer layer. A distinctive and widespread variety of patterned ground at high latitudes may not be related to the formation of an active layer. A small-scale polygonal pattern (dubbed basketball terrain by Malin and Edgett, 2001; Fig. 2) was probably formed as sublimation polygons that do not require an active layer, but instead require thermal oscillation at temperatures constantly below 0 1C (Marchant et al., 2002). On Mars, the year-average surface temperature is well below the water freezing point (0 1C) everywhere on the planet and the resulting permafrost covers all the planet and is currently several kilometers thick (e.g., Clifford, 1993, and references therein). The surface temperature on Mars under the present climate conditions routinely exceeds 0 1C in some places during some time of the day. This occurs in the equatorial zone throughout the whole year, and in mid-latitudes in summer, especially on equator-facing slopes. Thawing or sublimation of ice, if it exists at or very near the surface, can potentially occur under these transient conditions. Current knowledge suggests that there is no ice available for melting in the immediate vicinity of the surface in these regions (e.g., Mellon and Jakosky, 1993; Jakosky et al., 1993, 1995). Recent calculations show that if ice/snow/frost is present, true melting of small amounts of ice (rather than sublimation) will probably occur only under favorable local conditions (Hecht, 2002). Such melting, however, would not initiate process characteristic of an active layer because the maximum possible amount of liquid water is tiny (not more than a few millimeters) and exists for only a very short time (hours). Due to the high evaporation rate on Mars a diurnal freeze thaw cycle cannot currently lead to active layer processes typical of moist soils on Earth. Under different conditions on Mars in the past or in the future, active layers could potentially form due to seasonal thermal (thaw) waves, which imply positive temperatures (above freezing) to a depth comparable to the yearly thermal skin thickness, at least tens of centimeters. This requires that the day-average ( day-average means average over the whole spin period of the planet) surface temperature exceeds 0 1C during a warm season. Presently the day-average surface temperature is below 0 1C everywhere on Mars throughout the entire year (Fig. 3); therefore, formation of an active layer as defined here does not occur. In addition to the presence of permafrost and positive warm-season temperatures, active layer processes that produce geomorphic change require the availability of water that can freeze and melt. On the Earth, water is present virtually everywhere in the permafrost regions. On Mars, theoretical considerations (Mellon and Jakosky, 1993) predict ground ice beneath a thin dry soil layer at high latitudes and beneath a much thicker dry layer at low latitudes. Recent Odyssey GRS experiment data generally confirmed the current distribution of near-surface ground ice predicted by these models (e.g., Boynton et al., 2002; Feldman et al., 2002; Mitrofanov et al., 2002). Longer term insolation variations due to evolving spin and orbital parameters cause the latitudinal migration of near-surface water ice stability (Mellon and Jakosky, 1993, 1995; Mellon et al., 2004). Since there is presently no active layer on Mars, the numerous geological features characteristic of active layer processes are likely to have formed in the geological past under different climate conditions. The climate history of Mars is an exciting and controversial topic. In earliest Mars history (the Noachian), valley networks have been cited to support a warm and wet climate period during which conditions may have permitted pluvial activity (e.g., Jakosky and Phillips, 2001; Carr, 1996). In the Hesperian Period, the emplacement of large aqueous outflow channels following cryospheric breaching and groundwater discharge may have created regional standing bodies of water that in themselves influenced global atmospheric and climate conditions (e.g., Baker et al., 1991; Gulick et al., 1997). During the youngest Amazonian Period, postulated ice ages affecting mid-latitudes (e.g., Head et al., 2003) and tropical and mid-latitude mountain glaciers (e.g., Lucchitta, 1981; Head and Marchant, 2003; Shean et al., 2005; Head et al., 2005) provide evidence for large fluctuations in climatic conditions. These factors, together with recent studies of spin/orbital parameters predicting that the current insolation conditions are not typical

292 ARTICLE IN PRESS M.A. Kreslavsky et al. / Planetary and Space Science 56 (2008) 289 302 Fig. 3. Year-maximum day-average surface temperature on Mars. Surface temperature data are taken from the European Martian Climate Database available at http://www-mars.lmd.jussieu.fr/. This data set has been generated with a GCM (Forget et al., 1999), but represents well the observed Martian climate system. (e.g., Laskar et al., 2004), suggest the possibility that an active layer has been present on Mars in its past history and may have had an important influence on the geological and geomorphological record. Under what conditions will an active layer form on Mars, and when in the past history of Mars might this have occurred? We approach the problem of predicting the presence of an active layer by starting with the present, where conditions are most well known. We then examine climate conditions different from the present climate that are necessary to change conditions sufficiently to produce an active layer. Then we examine the statistics and distribution of steep slopes on Mars, which are most sensitive to change during periods of active layer formation. On the basis of these data, we infer the distribution and onset of active layer formation. Finally, we present evidence for the plausible timing of the most recent possible episodes of active layer formation and when it might occur again in the future. 2. Climate conditions necessary for active layer formation Astronomical forcing is the term used to describe the change of spin and orbit parameters that strongly influence the climate system of Mars (e.g., see the review by Kieffer and Zent, 1992). Climate change due to astronomical forcing is the most probable cause for the appearance and disappearance of an active layer. The orbital eccentricity of Mars and the position of its spin axis in space change with time due to gravitational perturbations from other planets. These changes cause fluctuations in insolation patterns and in this manner strongly influence the surface temperature and the behavior of volatiles and the atmosphere surface volatile exchange system. The insolation pattern is completely defined by three spin-orbital parameters: orbit eccentricity e, obliquity y (the angle between the orbital plane and the equatorial plane), and season of perihelion quantified as the areocentric longitude of perihelion from the moving equinox L P. Although all three parameters are essential in determining the characteristics of the climate system, obliquity has long been recognized as the most powerful climate driver (e.g., Kieffer and Zent, 1992). Obliquity oscillates with a period of 125 ka, and has an amplitude modulation with a period of 1.3 Ma. The maximum oscillation amplitude is 201. In addition to these quasi-periodic oscillations, there are chaotic changes of mean obliquity that occur at time scales of 5 Ma and greater (Laskar and Robutel, 1993). Due to the chaotic nature of these oscillations, and of solar system dynamics at longer time scales, predictive calculations of the spin/ orbital parameters are possible only back to 10 20 Ma before the present. The calculations show that the presentday oscillations around 251 obliquity were preceded by a period of generally higher obliquity (357101) prior to about 5 Ma ago. Recently, Laskar et al. (2004) performed an extensive series of calculations of spin/orbital parameters for much longer (250 Ma) time spans. The calculation runs differed due to tiny variations of the assumed moment of inertia of the planet. These calculations are not predictive for these time scales for the reasons described above, but they do provide a plausible picture of obliquity variation patterns that are likely to have existed in the Late Amazonian (Fig. 4). During this period, the obliquity is predicted to have varied from almost 01 to 651 with the mean value over the range of all calculations of 351. Of course, this mean value should not be considered as an estimate of the actual mean obliquity for the last 250 Ma, because the mean obliquity varies widely from run to run (Fig. 4). This mean value illustrates well, however, the fact that the present-day obliquity (251) is almost certainly moderately low compared to that of the recent geological past. Spin/orbital parameters control the surface temperature through a very complex climate system. The whole system

ARTICLE IN PRESS M.A. Kreslavsky et al. / Planetary and Space Science 56 (2008) 289 302 293 Fig. 4. Selected possible scenarios for the history of obliquity during the last 250 million years. Fifteen (15) of 1001 runs taken from Laskar et al. (2004). The obliquity history for the last 20 Ma is very similar for all 15 plots, strong differences for earlier epochs appear in calculations due to tiny differences in the assumed moment of inertia of Mars and are caused by chaotic nature of the solar system dynamics. Green line marks 451 obliquity considered as a rough proxy for active layer formation threshold. can be modeled with global climate models (GCMs) (Forget et al., 1999; Haberle et al., 2003; Richardson et al., 2003; Mischna et al., 2003), which reproduce well the present-day climate. Recently a number of studies have analyzed the behavior of the GCMs under higher obliquity (e.g., Richardson and Wilson, 2002; Mischna et al., 2003; Levrard et al., 2004). Since the models have a large number of adjustable parameters that are not fully constrained, the GCM predictions for different past climate conditions should be applied with these uncertainties in mind. The most uncertain GCM parameter is the atmospheric pressure at high obliquity. First, it is not clear how much CO 2 was available for atmospheric cycling during past periods of high obliquity (for example, an unknown portion of the CO 2 inventory may currently be stored within the polar layered deposits). Second, it is not clear whether additional available CO 2 would cause selfsustained greenhouse warming, or permanent solid CO 2 deposits would form instead. In this paper we consider the climate system qualitatively, rather than using specific climate models; with this approach we inevitably lose some of the accuracy of the more complex models, but we gain the ability to provide an overview of the principal trends regardless of specific model parameter variations. In this manner we can also make general predictions that can be further tested with more refined GCMs. The response of summer day-average surface temperature to variations in insolation patterns depends strongly on the atmospheric pressure. For example, for a dense atmosphere, such as that on the Earth, there is a strong thermal coupling between the atmosphere and the surface; the day-average surface temperature is close to the dayaverage air temperature. Therefore, for a given latitude, there is a strong elevation-related zonality in surface temperature caused by atmospheric lapse, and surface slopes have little effect on the surface temperature. On the other hand, for a thin atmosphere, such as that on presentday Mars, the surface temperature is primarily controlled by insolation. Therefore, there is no elevation-related temperature trend for a given latitude, and topography influences the temperature as long as it influences insolation. During periods of very low obliquity (o10 151), the atmosphere may completely collapse, because almost all atmospheric CO 2 condenses in the coldest areas at high latitudes (Kreslavsky and Head, 2005). Unlike the full global climate system, insolation itself can be accurately and readily calculated, especially if one assumes a transparent atmosphere. If we take surface slopes into account, the insolation has a complex dependence on season, latitude, and slope orientation reflecting a complex interplay between how high the sun rises and how much time it spends above the local horizon; several illustrative examples are shown in Fig. 5. Generally, as obliquity increases, the yearly insolation is redistributed from the equatorial zone to the polar regions. At the equator, the insolation pattern varies weakly with obliquity changes. Seasonal changes are minor and the overall differences between slopes of different orientation are not great unless obliquity is extremely high (4601). At high obliquity, polar regions receive higher insolation during spring and summer seasons, while slope orientation has little effect. In mid-latitude zones the seasons are more strongly expressed at high obliquity, and the insolation regime of slopes depends strongly on slope orientation. The pole-facing slopes at higher obliquity (4351) are more well illuminated in the summer than equator-facing slopes. The general dependence of summer day-average insolation on obliquity, latitude, and north south slopes is schematically summarized in Fig. 6a. Analysis of the full range of variations illustrated shows that the greatest insolation is reached at high obliquity at high latitudes, as well as on pole-facing slopes at mid-latitudes. For a thin atmosphere, such as that on present-day Mars, these are the regions where we might expect summer day-average temperatures to exceed 0 1C, and therefore, for an active layer to form. Insolation is of course not directly mapped into surface temperatures. Generally, however, an increase in insolation produces an increase in temperature, and we will derive approximate estimates of the threshold of active layer formation in terms of insolation. Climate models can be used to quantify these considerations (e.g., Paige, 2002). Costard et al. (2002) presented the results of calculated summer day-average surface temperatures with a modified GCM from Forget et al. (1999) for a range of obliquities, latitudes, and surface slopes. From these results we determined that the 0 1C isotherm corresponds to an above-atmosphere, day-average insolation in the range of

294 ARTICLE IN PRESS M.A. Kreslavsky et al. / Planetary and Space Science 56 (2008) 289 302 Fig. 5. Year-maximum day-average insolation for horizontal surfaces (bold line) and 301-steep slopes of different orientations (thin lines: blue marked with N is for the north-facing slopes, red marked with S is for the south-facing slopes, and green not marked is for the east- or west-facing slopes). The insolation (horizontal axis) is scaled as parts of the solar constant at the mean Mars distance form the sun, and is plotted as a function of the latitude (vertical axis). The three plots correspond to three different spin/orbit configurations: (1) the present, (2) the most recent well-pronounced obliquity peak 0.64 Ma ago, (3) the maximal extent of the active layer 9.45 Ma ago. Fig. 6. (a) Dependence of insolation on obliquity, latitude and north south facing slopes, shown in a qualitative scheme. (b) The global trend of steep north- and south-facing slope occurrence, shown in a qualitative scheme. 0.68 0.75, where insolation is measured in parts of the solar constant at the mean Mars distance from the Sun. The model uses the present available CO 2 inventory. We also applied a second approach by using the publicly available results of the GCM by Forget et al. (1999) as a good proxy for the present observed Mars climate. These data clearly show (Fig. 3) that the surface temperature is not solely a function of insolation: under the same aboveatmosphere, day-average insolation, but for different areas and seasons, the day-average temperature varies within a range as wide as 20 1C. These variations are mostly caused by surface albedo. Despite these wide variations, the general global trend of increasing temperature with increase of insolation is clearly seen (compare the latitudinal insolation trend in Fig. 5(l), bold curve, and the map in Fig. 3). For regions and seasons of the highest day-average temperatures on Mars ( 60 to 25 1C) we determined the regression of the temperature on the aboveatmosphere insolation (for horizontal surfaces). Then we extrapolated the regression line to 0 1C and obtained an insolation of 0.62. This is, of course, an extrapolation and does not have a well-established physical basis, but it does not involve the use of any poorly defined parameters. Below we use geologic observations to make an independent estimate of the onset and presence of active layer formation. 3. Properties of an active layer on Mars How do active layer properties and processes on Mars in the past differ from those currently observed on the Earth? On the basis of general considerations we expect a number of significant differences. Unless there is a thick atmosphere and strong self-sustaining greenhouse effect on

ARTICLE IN PRESS M.A. Kreslavsky et al. / Planetary and Space Science 56 (2008) 289 302 295 Mars, the year-average temperature is low, well below 0 1C, even if in the summer the temperature is high and active layer forms. This causes the permafrost below the permafrost table to be cold in comparison to typical terrestrial cases. Cold permafrost would cause more intensive upward freezing during the cold season (Fig. 1c); this will reduce effects related to the confinement of liquid water between freezing fronts. For the same reason, taliks (Fig. 1d) are not likely to occur on Mars. A number of factors influence the thickness of an active layer. The greater distance from the Sun is balanced by higher sun and long summer days at high obliquity. The longer martian year causes a longer warm season, which favors deeper penetration of the thawing front. On the other hand, the much lower cold-season temperatures slow the propagation of the thawing wave, which favors a thinner active layer. Generally, we would expect the same order of magnitude (tens of centimeters) for the thickness of the active layer on Earth and Mars. Due to weak thermal coupling between the atmosphere and the surface on Mars, the surface albedo causes great variations in the summer temperatures, which, in turn, have a strong effect on the existence of an active layer and its thickness. Thus, we expect very wide lateral variations of active layer properties related to local conditions. The nature and intensity of active layer processes depend not only on surface temperature, but also on the availability of water ice to melt and freeze, and on its abundance. As shown recently by GCM calculations, obliquity changes cause migration of water ice between polar regions and mid- and/or low-latitudes (e.g., Mischna et al., 2003; Levrard et al., 2004; Forget et al., 2006). Thus under some conditions even surface ice can be available for summer melting. Mellon and Jakosky (1995) show that the ground ice stability zone expands from high latitudes to mid-latitudes with an increase in obliquity, and there are no published predictions of ground desiccation in the polar region. Thus, it is quite probable that at least some ice is always available for active layer processes at medium to high latitudes. Another consequence of the thin and generally cold atmosphere on Mars is high evaporation rates. The availability of water in the soil depends on re-supply of water through condensation and precipitation, which is thought to be strongly controlled by local microclimate conditions. This would also contribute to the lateral variability of active layer properties. 4. The occurrence of steep slopes and implications for active layer formation One of the major manifestations of the presence of an active layer is the modification and degradation of slopes and a consequent reduction in slope magnitude (e.g., cryoplanation) (Fig. 7). Initial steep slopes are produced by a variety of processes, such as tectonic scarps associated with faulting and caldera walls associated with volcanism. Fig. 7. Left: a typical crater (6.5 km diameter) at mid-latitudes (43.51S, 194.81E). Portion of MOC NA image M23/00404, north is approximately on the top; illumination is from the upper part of the upper-left. Boxes show locations of two enlargements presented on the right. Vertical line shows location of MOLA profile shown on the upper right. Southern, equator-facing crater wall is steep and shows bedrock outcrops and masswasted debris. Northern, pole-facing wall has been modified by formation of an active layer and transport of debris down onto the crater floor, smoothing the crater rim and modifying the crater interior. Dissected and partly eroded layered mantle is clearly seen everywhere except on the steepest slopes. Fresh impact craters provide widespread examples of landforms characterized by steep interior walls and rim crests (e.g., Pike, 1980; Garvin, 2005). Thus, global altimetry data sets can be very useful in locating landforms with the steepest slopes (e.g., Kreslavsky and Head, 1999) and areas that lack steep slopes (Kreslavsky and Head, 2003) (Fig. 8), and thus identifying candidate regions or locations that might have been shaped by active-layer slope degradation. We have statistically examined north south slopes on Mars at a baseline of 0.3 km. The slopes were derived from precise MOLA topographic profiles (Smith et al., 2001). We have found a strong latitudinal trend in the occurrence of steep slopes on Mars (Kreslavsky and Head, 2003). The frequency of steep (4201) slopes drops more than three orders of magnitude from equatorial to high-latitude regions (Figs. 8 and 9). The boundary between steep slopes that are preserved and those that are reduced occurs at lower latitudes for pole-facing slopes and at higher latitudes for equator-facing slopes (Fig. 9). This produces a strong asymmetry in steep slopes in the mid-latitude zone around 451 latitude in both hemispheres (Fig. 10). The number of steep pole-facing segments in MOLA profiles in

296 ARTICLE IN PRESS M.A. Kreslavsky et al. / Planetary and Space Science 56 (2008) 289 302 Fig. 8. Distribution of slopes steeper than 301. High concentration of steep slopes at Valles Marineris walls is clearly seen. Note the paucity of steep slopes at mid-to-high latitudes in both hemispheres. Fig. 9. Abundance of N-facing and S-facing slopes relative to that of the typical equatorial highlands plotted against latitude for Terra Cimmeria (180 2201W). The actual quantity plotted is the proportion of MOLA profile segments of proper steepness and slope direction within a narrow latitudinal zone normalized by the same proportion for 10 201S zone representing typical equatorial highland terrain. The different curves correspond to different ranges of slopes in degrees, as shown, black curves, N-facing slopes, gray curves, S-facing slopes. these zones is almost a factor of three smaller than the number of steep equator-facing slopes. We interpreted these observations (Kreslavsky and Head, 2003) in the following way. Among the geological processes producing steep slopes, impact cratering operates continuously, producing slopes near the angle of repose (30 351) everywhere on the planet. In the equatorial zone there are many craters with walls steeper than 301. The virtual absence of very steep slopes (4301) and the strong deficiency of steep slopes (4101) at high latitudes is due to the operation of efficient processes of steep-slope removal. The universal latitude trend of these distributions points to a climate-related slope removal mechanism, and the dependence on slope orientation suggests that insolation had played a major role. Slopes are removed from exactly the same places that experience the highest day-average Fig. 10. Asymmetry of differential slopes calculated within 150-km wide latitudinal zones in Terra Cimmeria (180 2201W) and plotted against latitude. The measure of asymmetry is A ¼ (N N N S )/(N N +N S ), where N N and N S are numbers of steep N- and S-facing segments in all MOLA profiles. Different curves correspond to different ranges of differential slopes in degrees, as shown. surface temperatures through obliquity variations, that is, all slopes at high latitudes and pole-facing slopes at midlatitudes (see Figs. 6a and b). We infer that active-layer processes are the major slope-degradation and removal mechanism. Depending on ambient conditions and ground ice abundance, the following active layer processes can operate and contribute to slope removal: (1) segregation of meltwater and slope erosion by transient streams; (2) erosion by wet debris flows, the mechanism proposed by Costard et al. (2002) to explain the recent gullies; (3) solifluction and gelifluction, the latter being downslope movement of wet material over a frozen substrate and the former operating without an icy substrate; (4) cryoturbation, responsible for slope degradation according to Perron et al. (2003); and (5) cryoclastic erosion, the disintegration of rocks by water freezing in cracks.

ARTICLE IN PRESS M.A. Kreslavsky et al. / Planetary and Space Science 56 (2008) 289 302 297 These slope degradation processes operate during repeated and geologically short obliquity peaks (each lasting several tens of thousands of years). These processes may be intense enough to effectively remove slopes over geologically short periods of time, leaving few steep slopes at high latitudes. We performed a detailed study of the topography and morphology of all 130 craters in the 10 25 km diameter range within the typical northern plains (Kreslavsky and Head, 2006). Even the freshest crater in this population has walls 25 281 steep, less steep than 30 351 that should be expected for a fresh crater. The entire crater population was formed during the whole of the Amazonian (Head et al., 2002; Tanaka et al., 2005). Detailed analysis of the slope frequency distribution in this population indicated that intensive steep-slope removal operated (perhaps in repeated short episodes) at least during the last 200 400 Ma (Kreslavsky and Head, 2006), if we assume the Hesperian/Amazonian boundary to be at 3.1 Ga (Hartmann and Neukum, 2001). The occurrence of steep slopes is not the only global latitudinal trend observed in the statistical characteristics of topography. Among others, the most pronounced trend is a high-latitude decrease in subkilometer-scale roughness (Kreslavsky and Head, 2000). This trend occurs at similar latitudes, as does the deficiency of steep slopes; the smoothing boundary is diffuse and located at slightly higher latitudes that the onset of removal of equator-facing slopes. This smoothing is manifested as a softened appearance of topographic features in some mediumresolution (hundreds of meters scale) images of the surface. This effect was first noted in Mariner 9 images (Soderblom et al., 1973) and was initially attributed to eolian mantling. Squyres and Carr (1986) studied the same phenomenon of terrain softening in Viking images and attributed it to creep of a near-surface ice-rich layer. We avoid the term terrain softening because in earlier work (Squyres and Carr, 1986; Jankowski and Squyres, 1992; Squyres et al., 1992) this term was applied to two very different phenomena: (1) the soft appearance of topographic features at high latitudes mentioned above, and (2) latitude-dependent trends in crater depth, concentric crater infill, etc. The former is caused by smoothing of topography with a characteristic vertical scale on the order of meter(s), while the latter is related to creep of a hundreds-of-meters-thick layer; hence, the mechanisms, even if they are related genetically, are very different. A global latitudinal trend of roughness (Kreslavsky and Head, 2000) is observed for typical surface slopes at subkilometer scales. At these scales the typical slopes are gentle, and the vertical scale associated with subkilometerbaseline slopes is on the order of a meter. Thus, the observed high-latitude smoothing is caused by meter-scale changes in topography. Subkilometer-scale steep slopes are associated with hundred(s)-of-meters topographic variations, and meter-scale changes of topography are unable to remove them. Thus, the observed smoothing at high latitudes is not due to the same process as the removal of steep slopes. These processes, however, may be related genetically. The presence of active layer processes mobilizes surface material and should lead to surface smoothing (cryoplanation) by the diffusion creep mechanisms discussed above. Certainly, the active layer contributes to the roughness trend observed. Morphological observations with high-resolution images, however, clearly indicate the deposition of a smooth icy mantle at high latitudes as a powerful smoothing mechanism (Kreslavsky and Head, 2002; Head et al., 2003). This smooth mantle with sublimation polygons on its surface covers all small-scale surface features and obscures the potential morphological signs of active layer processes almost everywhere at high latitudes. At least the upper layer(s) of the mantle were probably emplaced later than the episodes of active layer formation had occurred, as discussed below. Thus, the most recent active layer is not the only contributor to terrain smoothing. Additional morphological observations are necessary to understand the relative roles being played by (1) active layer processes and (2) emplacement of a thin, latitude-dependent mantle in the creation of small-scale topography smoothing at high latitudes. In addition to the active layer processes listed above, flowing ice at the ground surface and ice-assisted creep at depth can and perhaps did act as eroding agents. A wide variety of small-scale ice flow features are observed in association with slopes at mid-latitudes (e.g., Milliken et al., 2003; Head et al., 2005). Erosion by flowing ice does not require temperatures above the melting point and does not imply active layer formation. Fast flow of thin ice, above or below the ground surface, is only possible when the temperatures are near the melting point and slopes are of sufficient steepness. Insolation conditions for true melting and for bringing near-surface ice close to the melting point are generally similar. This means that active flow of thin sheets of surface and subsurface ice could show geographic trends similar to those of active layers, that is, all slopes at high latitudes and pole-facing slopes at midlatitudes. Thus, flowing ice could contribute to the observed removal of steep slopes. We believe, however, that this contribution is generally less important than active layer processes for the following reasons. First, the boundary of steep-slope preservation is very sharp and consistent through the globe. This appears to be more consistent with a sharp threshold caused by a phase transition (solid ice to liquid water) rather than with gradational acceleration of ice flow. Second, the steep slopes of old geological features in the equatorial zone are well preserved, even though traces of regionally distributed glacial features are observed in this zone (e.g., Head and Marchant, 2003; Head et al., 2005). We now use the observed zonality of the occurrences of steep slope (e.g., Figs. 8 10) to independently derive the onset insolation for active layer formation from these geological observations (Kreslavsky and Head, 2004). The low-latitude onset of slope asymmetry occurs at 351 latitude for 201 steep slopes and at 421 latitude for 101

298 ARTICLE IN PRESS M.A. Kreslavsky et al. / Planetary and Space Science 56 (2008) 289 302 steep slopes (Fig. 10). This suggests that the onset of active layer formation corresponds to the insolation that occurs on these steep pole-facing slopes at these latitudes under the most favorable but still repeatedly occurring spin/orbital conditions. Favorable conditions are: high eccentricity ( ¼ closer proximity of the sun at perihelion), high obliquity, and perihelion close to the solstice. This latter condition occurs regularly. The recent calculations of the spin/orbit parameters of Mars (Laskar et al., 2004) showed that for the time span corresponding to the last 1 Ga, the maximum frequent value of eccentricity is 0.12, and that for obliquity is 54.51. Using the approach outlined above, we calculated that these values correspond to an onset day-average insolation of 0.83. This value is higher than the insolation corresponding to the 0 1C isotherm from GCMs (see discussion above). This difference is not surprising, because intensive active layer processes require somewhat warmer conditions than 0 1C on the hottest day in the summer. One important implication of the observed slope asymmetry at mid-latitudes is that its formation requires a thin atmosphere such as that on Mars today. As we noted above, if the atmospheric pressure is high enough, thermal coupling occurs between the atmosphere and the surface, and slope orientations have little effect on active layer formation. The observed very strong steep-slope asymmetry suggests that even at 501 obliquity the atmospheric pressure was not high. Furthermore, if we accept the interpretation that periods of active layer formation occurred during the entire Amazonian, then the observed strict zonality of these slopes is a long-term feature of the planet and this strongly implies that polar wandering (e.g., Schultz and Lutz, 1988) did not occur during the last 3 Ga. 5. Timing of an active layer in the recent past How often and over what time duration might an active layer have formed in the Late Amazonian? Laskar et al. (2004) have presented an extensive series of calculations of plausible obliquity variation patterns over the last 250 Ma (Fig. 4) that are not explicitly predictive due to the chaotic changes of mean obliquity at long time scales. Nonetheless, these patterns provide a framework in which to consider the role of an active layer in each of the sample histories. For simplicity, if we assume a value of 451 obliquity to delineate periods in which an active layer would form (4451 obliquity) or would not be predicted to form (o451 obliquity) we can examine the family of patterns to assess candidate scenarios. Obliquity patterns in which an active layer would be predicted to occur for extensive periods (4150 Ma) over the majority of the last 250 Ma are illustrated in Fig. 4e, f, i, m. On the basis of these assumptions, all but a few of these patterns (Fig. 4g, k, l, o) would have produced a period of active layer formation in excess of 10 Ma, a time interval certainly sufficient to produce significant surface modification. If an active layer is likely to have occurred repeatedly during the Amazonian Period, as we argue above on the basis of the nature and distribution of the steepest slopes, but does not currently occur on Mars due to present climate conditions, when in the recent past were the last times that an active layer was likely to occur? We used calculations from Laskar et al. (2004) of the spin/orbital parameters for the last 10 Ma to find when and where the formation of an active layer could occur in the recent past. If we formally apply the threshold of 0.83 for the dayaverage insolation we obtain the following results (Fig. 11). During the last 10 Ma period the insolation has never exceeded the threshold value in the southern hemisphere. The maximum insolation in the southern hemisphere, however, was very close to our inferred onset during four obliquity peaks (A in Fig. 11), reaching 0.823 at the south pole 5.30 Ma ago. In the northern hemisphere, four peaks of summer insolation exceeded 0.83 for the period from 7.75 8.09 Ma ago (B in Fig. 11). The second peak (7.86 Ma ago) was the highest, at 0.866 maximal insolation. Five more peaks occurred from 9.12 9.57 Ma ago, the fourth one yielded an absolute summer-time insolation maximum (0.872) on Mars during the last 10 Ma. This happened 9.46 Ma ago. For this highest peak, the summer insolation at the pole exceeded the onset limit for a period of 17 ka (9000 summers). At the very maximum part of the period the insolation exceeded the onset value down to 741N for horizontal surfaces. For 101 steep pole-facing slopes this zone extended down to 631N. For 201 steep pole-facing slopes, the active layer is predicted to be in the 871N 521N zone. Of course, for the reasons we describe above, these numbers should not be viewed as firm specific determinations. Instead, they provide an assessment of how the temporal and spatial distribution of the occurrence of an active layer might appear. The formation of an active layer, especially on a slope, is strongly controlled by the local environment. For example, at mid-latitudes alcoves on steep walls can shield the surface from the cold sky during the nights and can raise the day-average surface temperatures. The same topographic features at high latitudes can cast shadows during the day and lower the temperature. If the surface at high latitude were covered with bright ice, the formation of an active layer would not occur even at 901 obliquity. The nature and intensity of active layer processes depend not only on surface temperature, but also on the availability of water ice to melt and freeze, and on its abundance. As shown recently by GCM calculations, obliquity changes cause migration of water ice between polar regions and mid- and/or low-latitudes (e.g., Mischna et al., 2003; Levrard et al., 2004; Forget et al., 2006). Thus under some conditions even surface ice can be available for summer melting. Mellon and Jakosky (1995) show that the ground ice stability zone expands from high latitudes to mid-latitudes with an increase in obliquity, and there are no published predictions of ground desiccation in the polar region. Thus, it is quite probable that at least some ice is always available for active layer processes at medium to high latitudes. The proximity to the sun in summer due to

ARTICLE IN PRESS M.A. Kreslavsky et al. / Planetary and Space Science 56 (2008) 289 302 299 Fig. 11. Upper panel, the history of obliquity and eccentricity of Mars for the last 10 Ma, data from Laskar et al. (2004). Lower panel, year-maximum day-average insolation at the poles of Mars calculated from spin/orbital evolution for the last 10 Ma. Thin horizontal line marks the estimate of the active layer threshold. (a) Four insolation peaks in the southern hemisphere when insolation approaches the estimated threshold. (b) (c) Four and five insolation peaks in the northern hemisphere, when the insolation exceeded the estimated threshold. Vertical lines 2 and 3 mark points that correspond to panels 2 and 3inFig. 5. the eccentric orbit also plays an important role in producing a high insolation level. For example, the highest insolation peak 9.45 Ma ago occurred at 441 obliquity, noticeably lower than the maximum over 10 Ma (481). Recently, detection and mapping of a latitude-dependent smoothing of high latitude topography (Kreslavsky and Head, 2000, 2002), together with evidence for a latitudinally distributed dissected mantle (Mustard et al., 2001) and related features (Milliken et al., 2003) have led to the hypothesis that these deposits were emplaced during an ice age characterized by high amplitudes of obliquity oscillations that took place from about 0.4 2.1 Ma ago (Head et al., 2003). Subsequent to this time, as the amplitude of the obliquity of Mars decreased, the deposit has undergone desiccation in the 301 to 501N S latitude range. Could an active layer have been formed during this time period? This deposit is characterized by a series of distinctive morphological features that could be cited as evidence for the presence of an active layer. The most widespread is a type of pattern ground that resembles polygonal ground. This distinctive mounded terrain has been called basketball terrain due to its close similarity to the texture of the surface of a basketball (e.g., Kreslavsky and Head, 2003). Most polygons associated with active layers are characterized by angular patterns of ice wedges, sometimes with raised rims. The basketball terrain, however, widespread in the regions from 601 north and south latitudes to the polar deposits, consists of polygonal mounded patterns. These are very similar in morphology and scale to sublimation polygons found in the hyperarid, Mars-like polar deserts of the Antarctic Dry Valleys (Marchant et al., 2002). These features form from thermal cycling 50 1C in the absence of an active layer and derive their rounded morphology from sublimation processes operating preferentially at the cracked polygon margins. These sublimation polygons form in the upland frozen zone of the Antarctic Dry Valleys (Marchant and Head, 2005), and are very distinct in morphology and structure from the ice wedge and sand wedge polygons forming in the coastal thaw zone and mixed zone of the Antarctic Dry Valleys where active layers are present (Marchant and Head, 2005). Thus, the basketball texture