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. Kreslavskya,b,, James W. Heada, David R. Marchantc
a
Department of Geological Sciences, Brown University, Providence, RI 02912, USA
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
b
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).
0032-0633/$ - see front matter r 2007 Elsevier Ltd. All rights reserved.
doi:10.1016/j.pss.2006.02.010
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
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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
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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 621N 221E. 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
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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
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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 LP. 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
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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
CO2 was available for atmospheric cycling during past
periods of high obliquity (for example, an unknown
portion of the CO2 inventory may currently be stored
within the polar layered deposits). Second, it is not clear
whether additional available CO2 would cause selfsustained greenhouse warming, or permanent solid CO2
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
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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 CO2 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
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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 CO2 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
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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.
295
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
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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.
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 (NNNS)/(NN+NS), where
NN and NS 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.
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
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.
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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
297
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
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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
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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
3 in Fig. 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
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on Mars could have formed by thermal cycling in the
absence of an active layer, as predicted by non-active layer
conditions in the last 5 Ma (Fig. 11). Other features related
to the latitude-dependent layer and attributed to formation
in association with the ice age and its aftermath (Milliken
et al., 2003; Head et al., 2003) include dissected terrain,
seen in the 301 to 501N and S latitude regions. This consists
of irregular pits and depressions in the mantle that have
been described as ‘‘cryokarst’’ and interpreted to be due to
eolian deflation and sublimation of the ice-rich layer during
the interglacial period when obliquity decreased and ice
began to migrate from the lower latitude portions of the
mantle, where it was now unstable, back to the polar
regions. Lack of evidence of melting associated with these
features suggested that they represented ‘‘cryokarst’’ rather
than ‘‘thermokarst’’, where melting is implied. Viscous flow
features and gullies were also observed concentrated in the
30–501N and S latitude bands and are thought to be due to
local microenvironments (largely impact crater interiors)
that accumulated greater-than-average accumulations of
ice and snow, which underwent localized flow and minor
melting during and subsequent to the recent ice age period.
Thus, none of the major features observed in association
with the recent ice ages and their aftermath require the
presence of an areally extensive active layer, an observation
that is consistent with the predictions of the lack of an
active layer in the last 5 Ma (Fig. 11).
Geologically recent intra-crater ice-wedge polygons are
observed in 50–751 latitude zones in both hemispheres
(Mangold, 2005). The low-latitude boundary of these zones
is at somewhat lower latitude than the extent of the
geologically recent (in the 5–10 Ma period; Fig. 11) active
layer predicted above. This minor difference is not
surprising, bearing in mind the general nature of our
estimates. Small-scale sublimation polygons are formed on
top of the larger ice-wedge polygons (Fig. 2) (Mangold,
2005), which is consistent with the predicted absence of the
active layer in the most recent epoch (in the period
0–5 Ma). It is not clear why the ice-wedge polygons appear
only in craters and have not been observed outside craters.
Insolation conditions in the center of a flat crater floor and
on the surrounding plains are the same. Elevation itself
could not play a major role, because similar polygons are in
both hemispheres at very different absolute elevation. The
local environment in the crater interiors may favor
accumulation of larger amounts of water (ice/snow) ready
for a seasonal thawing/freezing cycle, or water might
preferentially flow to crater interiors in previous active
zone periods. An alternative explanation could be related
to albedo control of the surface temperature. Migration of
water ice during times of obliquity change are predicted to
form high-latitude icy mantles (Head et al., 2003). As
previously discussed, traces of such mantles are observed in
morphology (Kreslavsky and Head, 2000; Mustard et al.,
2001) and predicted by atmospheric general circulation
models (Levrard et al., 2004). Such a freshly formed mantle
may have high-albedo ice at its surface, which would
protect the surface from heating and prevent the formation
of an active layer. Craters may also act as traps for windblown dust and migrating sand and their floors could have
a relatively low albedo and be susceptible to increased
heating. Finally, the apparent superposition of the basketball terrain interpreted to be sublimation polygons on the
ice-wedge polygons (Fig. 2) suggests that these patterns
were formed in different periods of recent geological
history and that the period of active layer formation
has been superceded by a period without active layer
formation.
The specific predictions concerning the formation and
evolution of an active layer in the past history of Mars
described in this paper can be tested with existing data and
the distinctive landforms can be distinguished from those
produced by other processes. High-resolution MOC image
data (e.g., Malin and Edgett, 2001) provide the ability to
distinguish landforms at several meters-scale resolution,
and thermal emission spectrometer (TES) data and
THEMIS data permit the mapping of the thermal inertia
properties at several resolutions. Mars Express highresolution stereo camera (HRSC) image and stereo data,
as well as OMEGA imaging spectrometer data, provide
key contributions. Features identified and characterized
with these data can be compared with the topography and
slope data obtained by the Mars Orbiter Laser Altimeter
(MOLA), to test these hypotheses. Furthermore, upcoming
orbital missions (e.g., Mars Reconnaissance Orbiter) will
provide even higher surface image resolution and supporting data, and landers (e.g., Phoenix) and future increased
surface mobility will permit in situ analysis of surface
conditions and landforms.
When will conditions conducive to active layer formation recur in the future? Projections by Laskar et al. (2004)
to 10 Ma into the future predict that obliquity will be at or
below the range typical of the past few million years, and
thus conditions are not likely to be conducive to active
layer formation in the foreseeable future.
Acknowledgments
Discussions with N. Mangold were extremely helpful.
The work was partly supported by NASA Grant NAG512286 (MK) and NASA Grant NNG04GJ99G (JH), which
are gratefully acknowledged. Thanks are extended to Anne
Cote for help in manuscript preparation.
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