Icarus 208 (2010) 86–100
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Icarus
journal homepage: www.elsevier.com/locate/icarus
Supraglacial and proglacial valleys on Amazonian Mars
Caleb I. Fassett a,*, James L. Dickson a, James W. Head a, Joseph S. Levy b, David R. Marchant c
a
Department of Geological Sciences, Brown University, 324 Brook Street, Box 1846, Providence, RI 02912, USA
Department of Geology, Portland State University, 1721 SW Broadway, Portland, OR 97201, USA
c
Department of Earth Sciences, Boston University, Boston, MA 02215, USA
b
a r t i c l e
i n f o
Article history:
Received 21 September 2009
Revised 24 February 2010
Accepted 28 February 2010
Available online 11 March 2010
Keywords:
Mars
Mars, surface
Mars, climate
Geological processes
a b s t r a c t
Abundant evidence exists for glaciation being an important geomorphic process in the mid-latitude
regions of both hemispheres of Mars, as well as in specific environments at near-equatorial latitudes,
such as along the western flanks of the major Tharsis volcanoes. Detailed analyses of glacial landforms
(lobate-debris aprons, lineated valley fill, concentric crater fill, viscous flow features) have suggested that
this glaciation was predominantly cold-based. This is consistent with the view that the Amazonian has
been continuously cold and dry, similar to conditions today. We present new data based on a survey
of images from the Context Camera (CTX) on the Mars Reconnaissance Orbiter that some of these glaciers
experienced limited surface melting, leading to the formation of small glaciofluvial valleys. Some of these
valleys show evidence for proglacial erosion (eroding the region immediately in front of or adjacent to a
glacier), while others are supraglacial (eroding a glacier’s surface). These valleys formed during the Amazonian, consistent with the inferred timing of glacial features based on both crater counts and stratigraphic constraints. The small scale of the features interpreted to be of glaciofluvial origin hindered
earlier recognition, although their scale is similar to glaciofluvial counterparts on Earth. These valleys
appear qualitatively different from valley networks formed in the Noachian, which can be much longer
and often formed integrated networks and large lakes. The valleys we describe here are also morphologically distinct from gullies, which are very recent fluvial landforms formed during the last several million
years and on much steeper slopes (20–30° for gullies versus 10° for the valleys we describe). These
small valleys represent a distinct class of fluvial features on the surface of Mars (glaciofluvial); their presence shows that the hydrology of Amazonian Mars is more diverse than previously thought.
Ó 2010 Elsevier Inc. All rights reserved.
1. Introduction
Valleys resulting from water erosion provide critical clues to the
distribution, abundance and state of H2O throughout the history of
Mars, and insight into martian climate history (Carr, 1996).
Although early in Mars history, the climate may have been ‘‘warm
and wet” (e.g., Craddock and Howard, 2002), the Amazonian climate appears to have been cold and hyperarid throughout, comparable in many ways to certain microclimate zones in the Antarctic
Dry Valleys (Marchant and Head, 2007). During the Amazonian,
observations and modeling suggest that water on the surface and
in the near-subsurface has mainly been exchanged between major
reservoirs at the polar caps, the regolith, and extensive glacial
deposits at low-to-mid-latitudes (Forget et al., 2006; Madeleine
et al., 2009).
In this paper, we review some of the morphologic evidence for
cold-based glaciation at low-to-mid-latitudes on Mars and then
introduce new evidence that calls for localized erosion from supra* Corresponding author.
E-mail address: Caleb_Fassett@brown.edu (C.I. Fassett).
0019-1035/$ - see front matter Ó 2010 Elsevier Inc. All rights reserved.
doi:10.1016/j.icarus.2010.02.021
glacial and proglacial meltwater (Fig. 1). As an introduction, we begin with a brief overview of supraglacial meltwater streams and
landforms on Earth. We refer to the erosional features as ‘valleys’
on Mars, instead of ‘channels’ as might be more typical on Earth,
because of uncertainty that the features ever experienced bankfull
conditions and the lack of observable channel bedforms (Mars
Channel Working Group, 1983; see also Carrivick and Russell,
2006).
1.1. Surface melting of terrestrial cold-based glaciers and applications
to Mars
Relative to meltwater production in association with wet-based
glaciers, surface melting of cold-based glacier ice is minimal due to
the very low ice temperatures and low sensible heat available for
melting (e.g., Knighton, 1981; Fountain et al., 1998; Dyke, 1993;
Skidmore and Sharp, 1999; Atkins and Dickinson, 2007; Swanger
et al., 2010). Under optimal conditions, small amounts of seasonal
melting may arise from preferred insolation geometries (e.g., Fountain et al., 1998) and from melting alongside solar-heated debris on
the surface of otherwise relatively clean, cold-based glacier ice
C.I. Fassett et al. / Icarus 208 (2010) 86–100
87
Fig. 1. Examples of features on Mars interpreted to be formed by liquid water (CTX image number): (a) valley networks at 10.6°E, 22.8°S (P21_009049_1580), (b) a gully on
a crater rim at 104.6°E, 47.9°S (P18_008077_1317); and possible glaciofluvial features, (c) 11.7°E, 39.7°S (P15_006807_1391), (d) 164.4°E, 39°N (P17_007658_2175), (e)
58.3°E, 29°S (B01_010075_1498), and (f) 158°E, 42.8°S (P12_005877_1391). Locations for the glaciofluvial examples (and others in the paper) are shown on the distribution
map, Fig. 11.
(Shean et al., 2007). Two preferred insolation geometries include
terminal cliff faces (Fountain et al., 1998; Lewis et al., 1999) and
lateral margins where ice abuts steep bedrock slopes, the latter enable diffuse radiation to warm proximal ice surfaces. By way of
comparison, we note that meltwater channels in association with
wet-based glaciers typically incise thick deposits of proglacial outwash, and show anastomosing and/or braided patterns with multiple channels (Denton et al., 1993; Dyke, 1993). Channels from coldbased glaciers are typically shorter, fewer in number, and straighter (e.g., Dyke, 1993; Skidmore and Sharp, 1999; Atkins and Dickinson, 2007). Given the relative paucity of debris entrained in most
cold-based glaciers, lateral and proglacial channels may represent
the only evidence for cold-based glaciation (Kleman, 1994; Kleman
and Borgstrom, 1994; Atkins and Dickinson, 2007).
The key factors controlling the formation and evolution of surface meltwater and resultant streams from cold-based glaciers include: (1) elevation (temperature-dependence of melting;
Fountain et al., 1998, 2006; Lewis et al., 1998; Marchant and Head,
2007), (2) adjacent topography (controlling the focus and orientation of meltwater streams as well as the influx of solar radiation in
areas with considerable relief), (3) substrate type (till, bedrock,
polygonal ground, etc., controlling the amount and type of erosion
and pattern of streams; Atkins and Dickinson, 2007; Levy et al.,
2008), and (4) ice temperature (controlling the overall amount of
meltwater produced). The presence of supraglacial debris, if sufficiently thick, can serve to decouple underlying ice from atmospheric warming and prevent meltwater production (e.g.,
Kowalewski et al., 2006).
The main requirement for initiating top-down melting on Mars
is sufficient insolation to raise extant ice to its melting temperature, since the thin atmosphere has little heat content and minimal
effect on surface temperatures. One factor that may aid reaching
the melting temperature is if it is depressed by salts (e.g., Ingersoll,
1970; Knauth and Burt, 2002; Kreslavsky and Head, 2009). Sulfates
and other salts species are known to be common at the martian
surface (e.g., Clark and Van Hart, 1981; Rieder et al., 1997; Clark
et al., 2005), and may be transported and deposited onto ice by
the martian atmosphere (similar to atmospheric deposition of salts
in Antarctica today; Bao et al., 2000). Additionally, the factors listed
above for melting of terrestrial cold-based glaciers likely play a role
in enabling top-down melting of cold-based glaciers on Mars.
1.2. Low-to-mid-latitude cold-based glaciation on Mars
Observations from Mars orbit have increasingly led to the recognition of the importance of non-polar ice and glaciation at
low-to-mid-latitudes (e.g., Squyres, 1979; Lucchitta, 1981; Head
and Marchant, 2003, 2006; Milliken et al., 2003; Pierce and Crown,
2003; Kargel, 2004; Hauber et al., 2005, 2008; Head et al., 2005,
2006, 2009; Shean et al., 2005, 2007; Levy et al., 2007; Kadish
et al., 2008; Dickson et al., 2008; Fastook et al., 2009). Regions in
both the northern and southern hemispheres are characterized
by numerous lobate-debris aprons and lineated valley fill deposits,
interpreted as debris-covered glaciers (Head et al., 2006, 2009;
Head and Marchant, 2006) based on morphologic and topographic
evidence, and terrestrial analogs. Concentric crater fill in similar regions may also have a similar origin (e.g., Squyres and Carr, 1986;
Levy et al., 2009). Stratigraphy and crater counting suggest that
these features were last active during the Late Amazonian, with
crater retention ages younger than a few hundred million years
(e.g., Mangold, 2003; Head et al., 2005; Levy et al., 2009). Modeling
suggests that ice accumulation and glacial flow is favored at lowto-mid-latitudes during periods of higher obliquity (Forget et al.,
2006; Madeleine et al., 2009), and climate and glacial flow models
have successfully reproduced the locations and characteristics of
major areas of glaciation (Fastook et al., 2009) assuming specific
spin-axis/orbital configurations reasonable for Mars in the recent
past (Laskar et al., 2004). Terrestrial analogs in hyperarid, cold polar deserts, such as the Antarctic Dry Valleys, support the interpretation that these deposits on Mars were predominantly a result of
cold-based glaciation (Marchant and Head, 2007). Generally coldbased activity is inferred because features are almost pristinely
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preserved underneath glacial deposits (tills) and characteristic features such as drop moraines are commonly observed (Head and
Marchant, 2003).
The SHARAD radar sounder has recently provided new geophysical support for the geological interpretation that some of these
glacial features in both hemispheres remain ice-cored, covered
by only a thin, meters- to decameters-thick surface lag (Holt
et al., 2008; Plaut et al., 2009; Safaeinili et al., 2009). This is a
remarkable result considering that the crater populations on the
surface of these features imply crater retention ages of greater than
100 Myr, and subsurface ice is presently unstable against sublimation in the regions where they are found (e.g., Mellon and Jakosky,
1995). The reasons for this ice preservation are uncertain but may
reflect low long-term average net sublimation rates that led to
highly diminished rates of ice loss due to increasing sublimation
till thickness (Helbert et al., 2005), and/or long periods of higher
obliquity in the last 1 Gyr. Ice thicknesses were also much greater
during peak glacial conditions (Dickson et al., 2008; Marchant and
Head, 2008) and glacial ice may have been much more widespread
than that preserved in lobate-debris apron and lineated valley fill
deposits today (Head et al., 2009).
Despite the abundant evidence for ice emplacement, glaciation,
and the long-term survival of ice in these regions, there has been
little evidence of concomitant melting and runoff. Indeed, until recently, the most compelling evidence that has been presented for
glaciofluvial activity is a series of Hesperian age, braided ridges
in the south circum-polar Dorsa Argentea Formation (Head and
Pratt, 2001; Ghatan and Head, 2004) and large sinuous ridges in
Argyre Planitia (Kargel and Strom, 1992; Hiesinger and Head,
2002; Banks et al., 2009). Along with these eskers, long valleys also
emanate from the margins of the Dorsa Argentea Formation. These
are interpreted to represent Hesperian drainage of meltwater discharge from a south circum-polar ice sheet (Head and Pratt,
2001; Ghatan and Head, 2004). These features appear likely to
pre-date the widespread Amazonian ice deposits discussed above.
Other evidence for limited melting associated with glacial deposits
involves volcano–ice interactions (e.g., Chapman and Tanaka,
2002; Head and Wilson, 2007), in which subglacial and/or englacial
volcanism led to melting and the formation of small drainage valleys (Shean et al., 2005).
Most recently, Dickson et al. (2009) described evidence for
small valley-forms in a distinctive microenvironment on the floor
of Lyot crater in the northern mid-latitudes. The most plausible
sources of fluid for forming these valleys are ice-rich mantling
deposits on the crater floor or lobate-debris aprons, which are
probable debris-covered glaciers, along the crater walls. In the case
of Lyot, the observed valleys are known to be young (Amazonian),
both because of the youthful age of the Lyot crater itself (which
sets an absolute upper limit), as well as crater counts on the material the valleys incise (Dickson et al., 2009). The unique microenvironment of Lyot, which is a very deep crater and is thus among the
highest atmospheric pressure regions on Mars (Haberle et al.,
2001; Lobitz et al., 2001), has been interpreted to be a factor in
allowing liquid water to transiently exist there.
Here, we describe a broad survey of a distinct class of valleys
associated with features interpreted to be ice-related that we
interpret to be glaciofluvial (features whose origin is related to glacial meltwater). Fig. 1 shows these candidate features as well as
examples of valley networks (Fig. 1a) and gullies (Fig. 1b) for comparison. The key differences between the features we describe here
and classic Mars valley networks are the much smaller size of candidate glaciofluvial valleys, and their limited drainage development: valleys commonly have few or no tributaries and many
originate at a single source point. The glaciofluvial features we describe here also have clearly distinct morphologies from martian
gullies in that they lack alcoves and most lack fans (Malin and Edg-
ett, 2000), and occur on shallowly-sloping regions beneath ice-related landforms, rather than on the steep (20–30°) interior slopes
of crater walls.
The most common class of valley features we recognize occur at
the margins of materials interpreted to be of glacial origin and are
small (50–400 m wide) and short (<10 km in length) (Fig. 1c–e).
Our primary observational data are images from a survey of Mars
Reconnaissance Orbiter (MRO) Context Camera (CTX) data (Malin
et al., 2007), examined through the first eight PDS releases
(through mission part P22; 15,000 images were examined). For
specific features, we also incorporate data from MRO High Resolution Imaging Science Experiment (HiRISE) (McEwen et al., 2007a),
the Mars Odyssey Thermal Emission Imaging System (THEMIS)
(Christensen et al., 2003), the Mars Express High Resolution Stereo
Camera (HRSC) (Neukum et al., 2004), and the Mars Orbiter Laser
Altimeter (MOLA) (Smith et al., 1999).
The requirements for features to be included in our survey as
candidate glaciofluvial features are a direct inferred relationship
with current or past glacial deposits and characteristics suggesting
that they formed by the action of flowing water (based on aspect,
sinuosity, formation of branching networks, etc.). We exclude the
gully features (Malin and Edgett, 2000), which form on much steeper slopes and are morphologically distinct, although melting of
snowpack or ice is a possible origin for these features (Lee et al.,
2001; Costard et al., 2002; Christensen, 2003; Head et al., 2008;
Williams et al., 2009).
2. Observations
2.1. Examples of small valleys in and around craters
The most common setting in which we observe potential glaciofluvial valleys is on the interior and exterior of mid-latitude craters that are typically tens of kilometers in diameter. On Mars, the
geological setting and slope conditions that such mid-latitude (25–
55°) craters provide appear to have been preferred microclimates
for both ice accumulation (e.g., Squyres and Carr, 1986; Levy
et al., 2009) and melting (e.g., Costard et al., 2002).
2.1.1. Valleys within and outside an 80-km crater, 11.8°E, 40°S
Fig. 2 shows a series of small valleys on the interior of an unnamed 80-km crater. There are clear signs of past glacial ice on
the crater interior, particularly viscous flow features (e.g., Milliken
et al., 2003) that are perched high on the interior crater rim (elevation range 1100–2200 m), and large lobate flows on the crater
floor (elevation range 110 m to 500 m) (Fig. 2a and b). Stratigraphic relationships suggest that the viscous flow features are
younger than the broader lobes on the crater floor, implying multiple episodes of ice activity in this location (e.g., Head et al., 2008).
The potential glaciofluvial valleys begin at muted alcoves (Fig. 2c–
f) beneath viscous flow features at MOLA elevations of 150 m and
350 m. Thus, these valleys are interpreted to have formed during
earlier phases of glaciation, when ice was more widespread and
meltwater was available to incise the observed features. There
are also very small valleys on the crater exterior (Fig. 2b); since
these emerge from hollows filled with stipple-textured material,
similar to the texture of glacial remnants on the crater interior
and ice-rich material seen elsewhere (e.g., Mustard et al., 2001),
they may also have resulted from melting ice.
The valleys on the crater interior have near-constant widths and
few or no tributaries, consistent with being fed by point sources.
The longest valleys are 5 km in length, and the slopes of the observed valleys floors are between 2° and 6° (much lower than
slopes for gullies on Mars, 18–40°; e.g., Dickson et al., 2007; Parsons et al., 2008). At the termini of certain valleys (Fig. 2c–f), small
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Fig. 2. (a) Context image and mapping (b) of glacial units and small valleys on the interior of an 80-km, unnamed crater (11.8°E, 40.0°S). Glacial or ice-rich flow features on
the crater interior include viscous flow features (labeled vff) and a large lobate flow on the crater floor; white context box shows location of detail view. (c–f) Details of small
valleys with deformed fans, located beneath viscous flow features. Towards their fans, the valleys transition into ridges. (CTX images P13_005950_1401 and
P15_006807_1391, with HRSC nadir image 2694_0001 in the background.)
fans of material are observed, similar in morphology to fans in Lyot
crater (Dickson et al., 2009). The fan surfaces are rough-textured
and highly modified. Near their termini, valleys transition into
ridges, which we interpret to be due to post-fluvial terrain inversion (e.g., Williams and Edgett, 2005; Fassett and Head, 2007a;
Pain et al., 2007; Williams et al., 2007; Burr et al., 2009). Terrain
inversion may result from surrounding materials being removed
(by erosional processes, such as aeolian erosion or sublimation of
ice), leaving the coarser and more stable valley sediments as a
high-standing ridge.
On the basis of the measured Amazonian age of ice-related
deposits in these mid-latitude environments, valley formation in
this location is thought to have occurred at a similar time. However, a firm upper limit for valley formation comes from the age
of the host crater itself. We conducted crater counts on the host
crater and infer a Late Hesperian age for its formation based on
its superposed crater population, although the small surface area
for the ejecta imparts some uncertainty to this age (Fig. 3a). This
crater age is also supported by the fresh preservation state of the
crater and its ejecta, as well as the fact that secondary craters at
sizes of only a few hundred meters still remain recognizable on
surrounding plains. Thus, these valleys are clearly distinguished
in time from valley networks formed earlier, in the Noachian/Early
Hesperian on Mars (Fassett and Head, 2008a).
2.1.2. Valleys in a 70-km crater, 352.5°E, 41.5°S
Valleys are also found eroding material inside an unnamed 70km crater (Fig. 4a and b), with small exposures of concentric crater
fill material on the interior of the northern rim (Fig. 4c and d). The
host crater has an estimated crater retention age in the Late Hesperian or Early Amazonian (Fig. 3b) and a morphologically fresh
appearance consistent with this crater population. The concentric
crater fill we observe here has long been interpreted as ice-rich
(e.g., Squyres and Carr, 1986), an interpretation which has been
bolstered by recent spacecraft analyses which suggest that lobate-debris aprons, lineated valley fill, and concentric crater fill
all have a similar origin (e.g., Head et al., 2006, 2009; Levy et al.,
2009) and in some instances, have clear signs of extant ice (Holt
et al., 2008; Plaut et al., 2009; Safaeinili et al., 2009). Thus, we
interpret this fill material as a remnant glacial deposit. At present,
the accumulation zone directly adjacent to and south of the crater
rim (Fig. 4c) is generally free of the fill material, probably because
the glacier was beheaded (see Milkovich et al., 2006; Head et al.,
2008).
Although a few other valleys of <1 km in length are found inside
this crater, the only valley of substantial length is a 5.5 km long,
single valley that terminates in an elongate fan (Fig. 4e and f). This
valley has nearly constant width, a lack of tributaries and moderate
sinuosity, similar characteristics to valleys described in Section 2.1.1 and Fig. 2, which are at approximately the same latitude
and 850 km away. The average slope of the eroded valley floor is
5°. The valley headwaters were located at a re-entrant where a
remnant lobe remains today (Fig. 4e and f). We infer that ice was
advanced 1 km further into this re-entrant when melting occurred (where the valley originates) and has since retreated to its
present stand (Fig. 4f). This interpretation is consistent with previ-
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Fig. 4. (a and b) Context image and interpretative sketch of a fresh, 70-km
unnamed crater (352.5°E, 41.5°S); CTX image P16_007256_1383 and a THEMIS VIS
mosaic, superposed on a THEMIS IR daytime mosaic. (c and d) Detailed image and
interpretative sketch of the location for the observed small glaciofluvial valley,
emanating from probably ice-rich/glacial concentric crater fill. (e and f) Image and
sketch of the single 5.5 km long glaciofluvial valley in this crater, which
terminates in an elongate fan. The valley begins in a small alcove, where remnant
ice deposits are now 1 km from the valley head.
ous work on martian glacial landforms that has shown the glacial
remnants that exist today were preceded by a period when ice
was more extensive (Head and Marchant, 2003, 2009; Dickson
et al., 2008).
Fig. 3. (a) Crater count of the ejecta of the 80-km crater at 11.8°E, 40°S, yielding a
Late Hesperian age for the crater hosting ice and valley features in Fig. 2. (b)
Crater count of the ejecta of the 70-km crater at 352.5°E, 41.5°S, suggesting a Late
Hesperian or Early Amazonian age for this crater. (c) Crater count of the valleyincised surface beneath a series of viscous flow features in a crater at 113°E, 39°S.
There are very few craters on this surface, suggesting a Mid-Late Amazonian age
for the valleys. (d) Crater count of a lobate-debris apron surface in an Acheron
Fossae trough, whose surface has been incised by a small valley. The lobate-debris
apron has an inferred crater retention age of 80 Myr (in the Neukum system) or
110 Myr (in the Hartmann system), consistent with other lobate-debris aprons
on Mars. The valley incises the feature, so it must be of comparable age or younger.
[In the left column, data is plotted cumulatively and isochrons are from the
Neukum Production Function (see Ivanov, 2001); in the right column, data are
plotted in an incremental manner and isochrons are from Hartmann (2005).
The period boundaries used are those calculated in Fassett and Head (2008a)
and a best fit curve (in red) is calculated by minimizing the misfit of isochrons
to the data in a least-squares manner.] (For interpretation of the references to
color in this figure legend, the reader is referred to the web version of this
article.)
2.1.3. Valleys in a 75-km crater, 88°E, 27°S eroding a glacial moraine
The remnants of a probable debris-covered glacier bounded by
ridges (moraines) are observed on the interior northern wall of a
large crater in the southern highlands (Fig. 5). Topographic profiles
of the lineated terrain from MOLA suggest this remnant glacial feature is highly deflated in its interior, with a convex-down profile.
This clearly contrasts with well-preserved lobate-debris aprons
that have a convex-up shape (e.g., Li et al., 2005; Ostrach et al.,
2007). Thus, this feature may have lost more ice than many of
the other mid-latitude glacial features that have been described,
consistent with its location (27°) near the latitudinal limit (towards
the equator) of where pervasive ice-related features are found on
most of Mars (Head and Marchant, 2009).
Strikingly, a portion of the moraine that bounds the maximum
recent glacial extent has been breached by a 200-m wide valley
(Fig. 5c and d). This valley continues for 6 km with an average
slope of 6°, and at its terminus is a 1-km long sedimentary fan
(Fig. 5c and d).
The origin of this valley may be similar to valleys that erode terrestrial glacial moraines. Moraine-cutting streams can be initiated
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Fig. 5. (a and b) The interior northern wall of a 75-km crater in the southern highlands (88°E, 27°S) where remains of a feature interpreted to be a debris-covered glacier are
found, as well as a series of moraines, perched on material in the crater interior. CTX image P22_009797_1527. (c and d) A valley incises a portion of one moraine and
continues downslope for 6 km; at its terminus is a small sedimentary fan. There are other possible young valleys terminating in the crater center from the east.
either catastrophically, due to a flood triggered as a morainedammed lake overtops (e.g., Clague and Evans, 2000), or more
gradually, as the moraine is eroded or undercut (e.g., Swanger
et al., 2010). In this case, there are no signs of either a lake ever
having existed or of a flood origin for the observed valley, suggesting that the second, more gradual scenario is likely. The valleys
clearly incise the moraine as well as a part of the remnant glacial
surface. These constraints require that the valley is associated with
or post-dates the last period of major ice activity in the crater
(interpreted as Amazonian). Moreover, the ejecta and secondary
craters from the host crater are superposed on the Hesperian
ridged plains, implying that the host itself is Hesperian-aged or
younger. Thus, the best estimate for the age of valley formation
in this crater is Amazonian, as the observed valley cannot be older
than the Hesperian age for the crater itself.
2.1.4. Valleys in an ancient 45-km crater, 31.4°E, 31.7°S
At this locality, on the interior wall of a crater in the southern
highlands, a series of small sinuous valleys are seen below small lobate-debris aprons (Fig. 6). The characteristics of these valleys are
consistent with formation in direct association with glacial deposits. The valleys have nearly constant width (500 m), high sinuosity, and headwaters just below the present margin of the deposits
interpreted to be ice-related. Tributaries of the larger valleys are of
particular interest at this location, as they coincide with the largest, furthest advanced lobes of material interpreted to be of glacial
origin. Valleys descend approximately 600–750 m into the crater
over lengths of 8–10 km, with average slopes of 4°.
The constraints on the timing of these valleys are limited, since
the host crater for these valleys is degraded and interpreted to be
of Noachian age. However, the lobate features appear very fresh
and similar in morphology and freshness to other Late Amazonian
examples in this latitude range. In addition, the characteristics of
these valleys are similar to others described in this paper that
are found in association with lobate-debris aprons, and which are
interpreted to be due to melting of ice, overland flow, and fluvial
erosion. Thus, we hypothesize that these valleys are similar in
age to the other examples we describe, despite the lack of direct
constraints on the formation age for the observed valleys.
2.1.5. Valleys in 70-km crater, east of Hellas, 113°E, 39°S
The region east of Hellas is rich in ice-related landforms (Pierce
and Crown, 2003; Head et al., 2005) and Forget et al. (2006) have
proposed possible mechanisms for the region being an epicenter
of accumulation related to ice mobilized from the south polar
cap. Of particular interest are valleys found in association with
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Fig. 6. (a and b) Image and sketch showing the western interior of a 45-km degraded crater in the southern highlands (31.4°E, 31.7°S); concentric fill material is found at the
base of the wall. THEMIS VIS images V09811002, V15352007, V26958009. (c and d) Details of the valleys emanating from lobes of concentric crater fill on the crater interior
(CTX image P14_006503_1473).
and beneath viscous flow features in this area. A well-developed
example of such features is observed in the 70-km crater shown
in Fig. 7 (Berman et al., 2009). During the most recent period of glaciation, marked by the current extent of the viscous flow features
in this crater (Fig. 7), ice apparently did not extend to the headwater of the valleys in the center of the crater. However, crater counts
suggest that these viscous flow features are very young (perhaps
only 1–10 Myr; Arfstrom and Hartmann, 2005). During earlier
periods of glaciation, still probably dating to the Late Amazonian,
thick deposits of ice may have extended much farther downslope,
an interpretation supported by rough-textured fill material between the viscous flow features and the observed valleys, which
may be a remnant of this past glacial advance.
The dense, sub-parallel valleys seen here support the interpretation that they formed via ice-related melting, as they have a
poorly-integrated planform pattern, a very immature drainage system consistent with a transient melting mechanism. Direct timing
constraints on the formation of these valleys is difficult because of
the small surface area (and thus poor crater statistics), and the rapid degradation and destruction of small craters on Mars (and thus
the possible preferential loss of the few craters that do accumulate). However, there are very few superposed craters on these valleys, and none is larger than 400 m (in a count area of 60 km2).
Assuming that craters larger than 400 m superposed on the valleys
would survive from their time of origin, which is a reasonable
assumption given that the 200-m wide valleys remain sharp, this
implies that the valleys are younger than the Hesperian/Amazo-
nian boundary. The valleys may actually be significantly younger
(Middle to Late Amazonian) on the basis of the superposed crater-size frequency distribution relying on small craters (Fig. 3c).
2.2. Valleys associated with regional-scale ice deposits
Several lines of evidence suggest that some locations on Mars
have experienced widespread glaciation, including integrated flow
patterns in valley and trough systems in Acheron Fossae and Deuteronilus Mensae (Head et al., 2006; Head and Marchant, 2009),
flow between interconnected craters (e.g., the ‘‘Hourglass” craters;
Head et al., 2005), and broad glacial-like aprons (lobate-debris
aprons) on plains surfaces. As with deposits localized on the interior of craters, some of these systems appear to have melted at
their margins forming valleys. Here, we describe several examples
of glaciofluvial valleys associated with regional-scale ice deposits.
2.2.1. Acheron Fossae: Valleys on the surface of lineated valley fill,
230°E, 35.9°N
Lobate-debris aprons interpreted as debris-covered glaciers are
common in the troughs of Acheron Fossae (e.g., Dickson et al.,
2006a; Head and Marchant, 2009; Head et al., 2009) (Fig. 8). In
one of the fossae troughs, a single small sinuous valley is directly
superposed upon a lobate apron, traversing across its surface for
approximately 10 km, apparently parallel with flow lineations
(Fig. 8c and d). The surface it incises has a Late Amazonian crater
retention age, with best fit age estimates of 80 Myr (using the
C.I. Fassett et al. / Icarus 208 (2010) 86–100
93
Fig. 7. (a) Image and (b) sketch of numerous small, parallel valleys found beneath viscous flow features (113°E, 39°S) (CTX images P02_001964_1416 and P03_002320_1413).
At some point in the past, ice was probably more extensive, possibly extending to the source area of these valleys.
Neukum isochron system; Ivanov, 2001) or 110 Myr (in the Hartmann isochron system; Hartmann, 2005). Thus, the valley itself
must also be young (Late Amazonian) (Fig. 3d).
The valley begins near the north wall of the fossae, at the lobe
front (ice flow in this fossae was approximately from south to
north). It then continues across the debris apron surface, with a
direction consistent with the topographic gradient measured with
MOLA gridded topography; the valley is expressed in a series of
tight meanders (with typical wavelength 400 m). Presently, the
valley is less than 10 m deep and 100–150 m wide, widening
slightly downstream. Near the eastern margin of the lobe where
the valley ends (Fig. 8c and d), it transitions into a ridge, similar
to the positive sections in Fig. 2, although without a distinct depositional fan. The most probable process for inversion is that the valley sediments were preferentially preserved as ice downwasted
due to sublimation, resulting in inversion of relief. Alternatively,
the ridge may have formed by englacial or subglacial processes,
although evidence for wet-based behavior is otherwise lacking. A
small, degraded valley is also observed off the apron to the west
of this inverted ridge which may also be glaciofluvial in origin
(Fig. 8d).
A distinct albedo boundary along the northern wall of the fossae
may mark a past highstand of ice (Fig. 8b, dotted line). This feature
is continuously exposed along the wall for 40 km, consistent with
this explanation, suggesting that 100–200 m of ice may have been
lost from the lobate-debris apron.
2.2.2. Coloe Fossae: Valleys from lobate-debris aprons, 55.7°E, 39.9°N
The Coloe Fossae/Protonilus Mensae region along the northern
dichotomy boundary has been a site of extensive glaciation during
the Late Amazonian (Kargel, 2004; Dickson et al., 2006b, 2008;
Head et al., 2009). Evidence from past ice flow direction requires
that some valleys had ice thicknesses of at least 920 m (Dickson
et al., 2008). Near where the marker for this thick ice is found,
Dickson et al. (2006b) and Head et al. (2009) noted the existence
of a series of small valleys (1–7 km in length) trending away from
lobate-debris aprons from a trough wall. New CTX data (Fig. 9) provide an enhanced view of these valleys, which are morphologically
similar to valleys in Fig. 5 in the southern hemisphere. The stratigraphy here is complicated as the valleys have clearly been mantled
by recent material, perhaps from atmospheric deposition of ice
(stippled texture in Fig. 9c; see, e.g., Mustard et al., 2001), which
now appears similar to ‘brain terrain’ seen elsewhere on Mars
(Levy et al., 2009). Given this post-fluvial mantling, it is challenging
to directly constrain the timing of valley formation.
CTX and MOLA data suggest that all the valleys here start at
approximately the 2300 m elevation contour, 1 km from the
end of the present glacial lobe. The valleys begin at nearly their full
width (150–300 m), and have slopes of 1–3°. The elongate
depressions at the distal margins of the lobate-debris aprons that
separate the glacial remnant from the valleys may have resulted
from preferential loss of ice at the front of the glacier. Nonetheless,
the valleys localized nature, limited extent, and proximity to the
remnant glacial deposits support a glaciofluvial origin.
2.2.3. Eastern Hellas plateau: Melting of the ‘‘Hourglass”-shaped
glacier, 102.5°E, 39.1°S
One of the most striking features interpreted as a glacier on
Mars is a lobate-debris apron with a distinctive appearance adjacent to a 4 km high massif east of the Hellas basin (Head et al.,
2005) (Fig. 10). The head of the lobate-debris apron is in the upslope 10 km crater and flows through a narrow gap into an adjacent
16 km crater (Head et al., 2005). Fine flow lineations on the lobate
apron surface trace across the apron surface and constrict as they
pass through the gap (Head et al., 2005).
CTX data reveal that the rim on the lower (16-km) Hourglass
crater is incised or breached, and valleys are present south and
west of this rim (Fig. 10c and d). These valleys outside the rim
are up to 800 m wide and 12 km long, and erode the crater’s ejecta and surrounding plains; these valleys are most deeply incised
beginning 6 km away from the rim. Nearer to the rim is a fan-like
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C.I. Fassett et al. / Icarus 208 (2010) 86–100
Fig. 8. (a and b) Image and interpretative sketch of a large trough within the Acheron Fossae in northern Tharsis (230°E, 35.9°N). The trough is filled with lineated valley fill,
the surface of which is incised by a small valley. Downhill (the direction of glacial flow) was to the north (right in this image); CTX images P01_001590_2160,
P02_001933_2174, P03_002144_2165, P04_002632_2173, P05_003067_2160, and P15_006825_2179). (c and d) A small, degraded, and highly sinuous valley that incised the
lineated valley fill surface in this location. At the margin of the lineated valley fill, it transitions into a positive ridge (c).
sedimentary deposit that appears to emerge from the incised notch
or breach.
There are at least two plausible mechanisms consistent with the
observations of the incision of the 16-km Hourglass crater rim and
the valleys on its exterior; in either case, the source of the water
was likely top-down melting of the ice-rich lobe within the 16km crater. In the first scenario, the rim of the crater was an
impoundment to this meltwater, leading to ponding until the crater rim crest was breached. Rapid drainage and down-cutting
would then occur, which is consistent with the larger channels observed here than in other proglacial valleys, which may have re-
sulted from more gradual erosion (compare to examples in
Figs. 2–9). Alternatively, if thicker ice was present in the Hourglass
crater (>100 m above present surface), supraglacial streams may
have flowed to the glacier’s margins at the crater rim, incised the
rim notch, and eroded the valleys outside the crater. This second
mechanism requires significant downwasting of the glacier to
reach the present state; some observations supporting such downwasting (possible moraines and other glacial remnants) are present, particularly near the narrow gap of the Hourglass. Given
these two plausible formation scenarios, these valleys may have
resulted from either proglacial or supraglacial melting.
C.I. Fassett et al. / Icarus 208 (2010) 86–100
95
Fig. 9. (a and b) Lobate-debris aprons along a plateau margin in the Coloe Fossae region where small valleys are found at apron termini (55.7°E, 39.9°N). HRSC nadir image
h5299 and CTX image P16_007161_2209. (c and d) Close-up view of the valleys. The stippled ‘‘brain terrain” texture material (Levy et al., 2009) appears to be superposed
upon the valleys and thus to post-date fluvial activity.
The valleys around the Hourglass glacier incise a surface with a
Hesperian crater retention age (Head et al., 2005). The surface of
the Hourglass glacier itself is much younger (Late Amazonian), in
the same age range as other lobate-debris aprons (Head et al.,
2005). Thus, the best estimate for when the observed valleys
formed is during the Amazonian, when the Hourglass glacier must
have been thicker and more active, and post-dating the Hesperian
plains and the ejecta from the larger Hourglass crater. A few other
valleys besides those directly sourced from the gap in the Hourglass crater are also found on the surrounding plains. Although
these have an uncertain timing and origin, they may also relate
to melting of glacial ice during a more extensive past glacial phase,
earlier in the Amazonian.
2.3. Geographic and elevation distribution
We have assessed the probable geographic distribution of these
features by examining CTX images through mission phase P22.
Although the features we describe in detail in this paper are some
of the best examples of valleys related to the melting of Amazonian
glaciers on Mars, many other candidate features exist. The broader
geographic distribution of potential glaciofluvial valleys is shown
in Fig. 11.
The geographic distribution of the features conforms directly to
the regions of Mars where features interpreted as glacial in origin
are most prevalent (Squyres, 1979; Squyres and Carr, 1986; Millik-
en et al., 2003; Head and Marchant, 2009; Head et al., 2009). Indeed, almost all areas at low-to-mid-latitudes where ice appears
to have been concentrated in the past show at least some signs
of possible small glaciofluvial valleys. This broad distribution suggests that conditions that allowed transient melting were not
uncommon although there is a latitude dependence that is probably a direct function of where Amazonian glaciation occurred. The
geographic distribution in Fig. 11 conforms well to the latitudes
where gullies are also the most dense (Malin and Edgett, 2000),
although the features we map here are on lower slopes (<10°) than
gullies, which erode slopes of 15–35° and these candidate glaciofluvial landforms are also older.
Along with their geographic range, the elevation range for the
glaciofluvial features is similar to gullies. There are very few high
elevation features (>2000 m above the datum), despite the existence of several large glacial deposits at these elevations. The glaciers on the west flank of the Tharsis Montes (Head and
Marchant, 2003; Shean et al., 2005, 2007) and Olympus Mons
(Head et al., 2005; Milkovich et al., 2006) all lack evidence for
glaciofluvial valleys similar to the others we describe. At present,
the only candidate glacial valleys associated with these glaciers
appear to involve ice–volcano interaction (Shean et al., 2005;
Head and Wilson, 2007; Kadish et al., 2008). At these high
elevations, the low pressure may have posed a barrier to the initiation of melting under typical conditions despite available ice
reservoirs.
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C.I. Fassett et al. / Icarus 208 (2010) 86–100
Fig. 10. (a and b) View of the ‘‘Hourglass” glacier flowing between two craters in the Hellas Montes, originally described by Head et al. (2005) (102.5°E, 39.1°S). CTX images
P02_001938_1408, P03_002149_1409, and P16_007397_1382. (c and d) Valleys on the plains outside the Hourglass glacier, emanating from a gap in one of the Hourglass
crater walls.
Fig. 11. Distribution map of glaciofluvial features that we interpret to be the result of melting of near-surface ice. Locations for other figures in this paper are labeled, as are
major regions for mid-latitude ice emplacement and glacial activity.
C.I. Fassett et al. / Icarus 208 (2010) 86–100
97
3. Glaciofluvial valleys in the context of valley network
formation over time
The age of these glaciofluvial features is interesting because
most of the fluvial record on Mars appears to be older, forming
in the Noachian and early in the Hesperian (e.g., Pieri, 1980; Carr
and Clow, 1981; Fassett and Head, 2008a). There are well-documented exceptions to the older fluvial features, however, and
some Hesperian and Amazonian terrains on Mars show clear evidence for fluvial erosion. The most prominent examples are along
the rim of Valles Marineris (Mangold et al., 2004, 2008; Weitz
et al., 2008), the Valles Marineris interior (Quantin et al., 2005),
on certain volcanoes (Gulick and Baker, 1990), and in the vicinity
of some young craters (Mouginis-Mark, 1987; Brackenridge,
1993; Williams and Malin, 2008; Morgan and Head, 2009). Burr
et al. (2009) also interpret the sinuous ridges associated with
the Medusae Fossae Formation as young (late Hesperian to middle Amazonian), although constraining the age of fluvial activity
is challenging because the material in which they are found is
easy to erode and the ridges appear exhumed; an alternative
view is that the lower units of the Medusae Fossae Formation
may be relatively old (Early Hesperian or before; Kerber and
Head, 2009).
The existence of demonstrably Hesperian to Amazonian-aged
valley networks has led some workers to argue that the global climate conditions responsible for early valley network formation
(which may be warmer and wetter than today) lasted well into the
Hesperian or even Amazonian (e.g., Craddock and Howard, 2002).
Conversely, other researchers have suggested that the existence of
these younger valleys may mean that the Noachian/Early Hesperian
valley networks could have formed under a climate more similar to
the modern cold, hyperarid than is commonly inferred (Carr and
Head, 2003; McEwen et al., 2007b). Furthermore, it is likely that
many of these fluvial features, particularly those associated with volcanic edifices, do not represent fundamental changes in the atmosphere of Mars, but rather represent local conditions related to the
internal supply of heat to melt snow and ice accumulated on volcanoes during obliquity-driven climatic excursions (e.g., Fassett and
Head, 2006, 2007b).
Thus, it is important to place the previously known young valleys
and the possible glaciofluvial valleys described here in context in
terms of both age and their climate implications. Most of the crater
ages and stratigraphic data for these valleys suggest that they are
Amazonian in age. In a few instances, these constraints require an
Amazonian age, such as where a valley incised a glacial moraine in
a young crater and where a valley incised the surface of the lobatedebris apron in Acheron Fossae.
In terms of climate requirements, the glaciofluvial valleys we
observe are less integrated, and typically far smaller than mapped
ancient valley networks (Fig. 1). The ancient valleys on Mars also
existed in an environment which allowed for large (and widely distributed) lakes to exist on the surface (Fassett and Head, 2008b,
and references therein). The glaciofluvial valleys are also sparsely
distributed, though not uncommon, at latitudes >26° in each hemisphere (Fig. 11). These characteristics suggest they may not have
required stable liquid water; instead, perhaps their formation only
required transient metastablity (e.g., Hecht, 2002). Their direct
association with cold-climate features is also an argument for their
formation in a cold-climate similar to the martian climate today.
As has been long noted, most surfaces of Amazonian and Hesperian
age lack any indication of fluvial modification (e.g., Pieri, 1980;
Carr and Clow, 1981), so the ‘wet’ conditions that formed these features must be far more limited than the Noachian or earliest Hesperian valley systems. We thus discount the likelihood that
transient regional to global rainfall is a reasonable explanation
Fig. 12. A schematic representation of valley network intensity over martian
history, modified after Fassett and Head (2008a). Ancient, highland valley networks
date predominantly to the end of the Noachian or earliest Hesperian, and the last
widespread activity appears to have ended by the Early Hesperian. The long-term
intensity of valley network activity in the earlier Noachian is not known. There are
several regions that experienced punctuated valley formation after the end of this
early period, such as on the plateau above Echus Chasma, on the volcanoes Hecates
Tholus, Ceraunius Tholus, and Alba Patera, and around several young craters (not
shown). (Arrows are meant to give a sense that the actual age is unknown; age
estimates for these young valleys generally overlap with each other based on formal
statistics, but are not consistent with Noachian formation; see detailed crater
counts in Fassett and Head (2008a).) The age of valleys in Lyot crater are estimated
to be younger than 800 Myr in the Hartmann system or 1.5 Gyr in the Neukum
system based on crater counting of the unit they incise (Dickson et al., 2009). The
valley cutting across the Acheron Fossae lobate-debris aprons (Fig. 8) are also
constrained by the age of the apron to be younger than approximately 100 Myr
(best estimates for the age of the apron is 80 Myr in the Neukum system or
110 Myr in the Hartmann system). The evidence we present in this paper requires
the existence of small scale, low-intensity valley formation at some points during
the Amazonian as well – perhaps in a periodic manner during or following peak
periods of mid-latitude glacial activity.
for the observed valleys. Instead, the source of water for forming
these valleys is almost certainly the nearby available inventory of
glacial ice.
Thus, under some conditions in the Amazonian, melting of surface or near-surface ice must have been possible. The local nature
and morphological distinctiveness of these small, glaciofluvial valleys (and other young valley systems) continue to support the
view that Noachian valleys formed in a different climate from
what characterized Mars during the Amazonian. In Fig. 12, we
present an updated schematic diagram for the history of valley
formation on Mars, illustrating changes in erosion intensity over
time. We emphasize that the glaciofluvial valleys we describe
here, as well as the other examples of young valley networks on
Mars, appear to be qualitatively different from Noachian valley
networks and very recent Amazonian gullies, and represent a distinctly different type of fluvial activity. In sum, the diverse record
of valley formation on Mars shows a long history of surface erosion under a range of conditions, and new spacecraft data will
undoubtedly lead to a greater understanding and appreciation
of this record.
4. Conclusion
We document the existence of small, glaciofluvial valleys associated with major Amazonian ice deposits in the mid-latitudes of
Mars. The meltwater production that formed these valleys may
be due to anomalous insolation conditions in climatic microenvironments. The formation of these features demonstrates the diversity of geomorphic processes that have occurred on Mars, despite
what remains strong evidence for long-term cold and dry condi-
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C.I. Fassett et al. / Icarus 208 (2010) 86–100
tions during the Amazonian. Even when Mars has primarily been
an icy world rather than a wet one, localized flow of water has occurred at its surface.
Acknowledgments
We thank David Baker, Seth Kadish, Ailish Kress, Gareth Morgan, and Sam Schon for helpful discussions and John Huffman for
assistance in stereo analysis and visualization. This work was supported in part by the Mars Data Analysis Program (MDAP
NNGO4GJ99G), the Mars Fundamental Research Program (MFRP
GC196412NGA; NNX06AE32G), the Applied Information Systems
Research Program (AISR NNXO8AC63) and the Mars Express High
Resolution Stereo Camera investigation (JPL 1237163), which are
gratefully acknowledged. We also acknowledge the efforts of the
MRO CTX science and engineering teams for obtaining data without which this study would not have been possible.
References
Arfstrom, J., Hartmann, W.K., 2005. Martian flow features, moraine-like ridges, and
gullies: Terrestrial analogs and interrelationships. Icarus 174, 321–335.
Atkins, C., Dickinson, W.W., 2007. Landscape modification by meltwater channels at
the margins of cold-based glaciers, Dry Valleys, Antarctica. Boreas 36, 47–55.
Banks, M.E., Lang, N.P., Kargel, J.S., McEwen, A.S., Baker, V.R., Grant, J.A., Pelletier,
J.D., Strom, R.G., 2009. An analysis of sinuous ridges in the southern Argyre
Planitia, Mars using HiRISE and CTX images and MOLA data. J. Geophys. Res., in
press.
Bao, H., Campbell, D.A., Bockheim, J.G., Theimans, M.H., 2000. Origins of sulphate in
Antarctic dry-valley soils as deduced from anomalous 17O compositions. Nature
407, 499–502.
Berman, D.C., Crown, D.A., Bleamaster, L.F., 2009. Degradation of mid-latitude
craters on Mars. Icarus 200, 77–95.
Brackenridge, G.R., 1993. Ancient martian valley genesis and paleoclimatic
inference: The present as a key to the past. Workshop on Early Mars: How
warm and how wet? LPI Technical Report, 93-03, pt. 1, pp. 2–3.
Burr, D.M., Enga, M.-T., Williams, R.M.E., Zimbleman, J.R., Howard, A.D., Brennard,
T.A., 2009. Pervasive aqueous paleoflow features in the Aeolis/Zephyria Plana
region, Mars. Icarus, 200.
Carr, M.H., 1996. Water on Mars. Oxford University Press, New York. 229 pp.
Carr, M.H., Clow, G.D., 1981. Martian channels and valleys; their characteristics
distribution, and age. Icarus 48, 91–117.
Carr, M.H., Head, J.W., 2003. Basal melting of snow on early Mars: A possible origin
of some valley networks. Geophys. Res. Lett. 30, 2245. doi:10.1029/
2003GL018575.
Carrivick, J.L., Russell, A.J., 2006. Glacial landforms, sediments: Glacifluvial
landforms of deposition. In: Elias, S.A. (Ed.), Encyclopedia of Quaternary
Science, pp. 909–920.
Chapman, M.G., Tanaka, K.L., 2002. Related magma–ice interactions: Possible
origins of chasmata, chaos, and surface materials in Xanthe, Margaritifer, and
Meridiani Terrae, Mars. Icarus 155, 324–339.
Christensen, P.R., 2003. Formation of recent martian gullies through melting of
extensive water-rich snow deposits. Nature 422, 45–48. doi:10.1038/
nature01436.
Christensen, P.R.and 10 colleagues, 2003. The Thermal Emission Imaging System
(THEMIS) for the Mars 2001 Odyssey Mission. Space Sci. Rev. 110, 85–130.
Clague, J.J., Evans, S.G., 2000. A review of catastrophic drainage of moraine-dammed
lakes in British Columbia. Quarternary Sci. Rev. 19, 1763–1783. doi:10.1016/
S0277-3791(00)00090-1.
Clark, B.C., Van Hart, D.C., 1981. The salts of Mars. Icarus 45, 370–378.
Clark, B.C.and 22 colleagues, 2005. Chemistry and mineralogy of outcrops at
Meridiani Planum. Earth Planet. Sci. Lett. 240, 73–94.
Costard, F., Forget, F., Mangold, N., Peulvast, J.P., 2002. Formation of recent martian
debris flows by melting of near-surface ground ice at high obliquity. Science
295, 110–113. doi:10.1126/science.1066698.
Craddock, R.A., Howard, A.D., 2002. The case for rainfall on a warm, wet early Mars.
J. Geophys. Res. 107, 5111. doi:10.1029/2001JE001505.
Denton, G.H., Sugden, D.E., Marchant, D.R., Hall, B.L., Wilch, T.I., 1993. East Antarctic
ice sheet sensitivity to Pliocene climatic change from a Dry Valleys perspective.
Geogr. Ann. A 75, 155–204.
Dickson, J.L., Head, J.W., Kreslavsky, M.A., Marchant, D.R., 2006a. Linear lobate
debris aprons, piedmont-like lobes, and crater fill in the Acheron graben region,
Mars: Evidence for debris-covered glacier formation and flow. Lunar Planet. Sci.
37. Abstract 1321.
Dickson, J.L., Head, J.W., Marchant, D.R., 2006b. Modification of graben along the
dichotomy boundary in eastern Arabia Terra (Coloe Fossae; 53–59°E, 37–41°N):
Nature and evolution of lobate debris aprons and their relationships to lineated
valley fill and glacial systems. Lunar Planet. Sci. 37. Abstract 1317.
Dickson, J.L., Head, J.W., Kreslavsky, M.A., 2007. Martian gullies in the southern midlatitudes of Mars: Evidence for climate-controlled formation of young fluvial
features based upon local and global topography. Icarus 188, 315–323.
Dickson, J.L., Head, J.W., Marchant, D.R., 2008. Late Amazonian glaciation at the
dichotomy boundary on Mars: Evidence for glacial thickness maxima and
multiple glacial phases. Geology 36, 411–414.
Dickson, J.L., Fassett, C.I., Head, J.W., 2009. Amazonian-aged fluvial valley systems in
a climatic microenvironment on Mars: Melting of ice deposits on the interior of
Lyot Crater. Geophys. Res. Lett. 36, L08021. doi:10.1029/2009GL037472.
Dyke, A.S., 1993. Landscapes of cold-centred Late Wisconsinan ice caps, Arctic
Canada. Prog. Phys. Geogr. 17, 223–247.
Fassett, C.I., Head, J.W., 2006. Valleys on Hecates Tholus, Mars: Origin by basal
melting of summit snowpack. Planet. Space Sci. 54, 370–378.
Fassett, C.I., Head, J.W., 2007a. Layered mantling deposits in northeast Arabia Terra,
Mars: Noachian–Hesperian sedimentation, erosion, and terrain inversion. J.
Geophys. Res. 112, E08002. doi:10.1029/2006JE002875.
Fassett, C.I., Head, J.W., 2007b. Valley formation on martian volcanoes in the
Hesperian: Evidence for melting of summit snowpack, caldera lake formation,
drainage and erosion on Ceraunius Tholus. Icarus 189, 118–135.
Fassett, C.I., Head, J.W., 2008a. The timing of martian valley network activity:
Constraints from buffered crater counting. Icarus 195, 61–89.
Fassett, C.I., Head, J.W., 2008b. Valley network-fed, open-basin lakes on Mars:
Distribution and implications for Noachian surface and subsurface hydrology.
Icarus 198, 37–56.
Fastook, J.L., Head, J.W., Madeline, J.-B., Forget, F., Marchant, D.R., 2009. Modeling
northern mid-latitude glaciation with GCM-driven climate: Focus on
Deuteronilus–Protonilus Mensae valleys. Lunar Planet. Sci. 40. Abstract 1144.
Forget, F., Haberle, R.M., Montmessin, F., Levrard, B., Head, J.W., 2006. Formation of
glaciers on Mars by atmospheric precipitation at high obliquity. Science 311,
368–371.
Fountain, A.G., Dana, G.L., Lewis, K.J., Vaughn, B.H., McKnight, D., 1998. Glaciers of
the McMurdo Dry Valleys, Southern Victoria Land, Antarctica. In: Priscu, J.C.
(Ed.), Ecosystem Dynamics in a Polar Desert: The McMurdo Dry Valleys,
Antarctica, pp. 65–76. In: American Geophysical Union Antarctic Research
Series, vol. 72. Washington, DC, 369 pp.
Fountain, A.G., Nylen, T.H., MacClune, K.L., Dana, G.L., 2006. Glacier mass balances
(1993–2001), Taylor Valley, McMurdo Dry Valleys, Antarctica. J. Glaciology 52,
451–462.
Ghatan, G.J., Head, J.W., 2004. Regional drainage of meltwater beneath a Hesperianaged south circumpolar ice sheet on Mars. J. Geophys. Res. 109, E07006.
doi:10.1029/2003JE002196.
Gulick, V.C., Baker, V.R., 1990. Origin and evolution of valleys on martian volcanoes,
J. Geophys. Res. 95, 14325–14344.
Haberle, R.M., McKay, C.P., Cabrol, N.A., Grin, E.A., Schaeffer, J., Zent, A.P., Quinn, R.,
2001. On the possibility of liquid water on present day Mars. J. Geophys. Res.
106, 23317–23326.
Hartmann, W.K., 2005. Martian cratering 8: Isochron refinement and the
chronology of Mars. Icarus 174, 294–320.
Hauber, E., van Gasselt, S., Ivanov, B., Werner, S., Head, J.W., Neukum, G., Jaumann,
R., Greeley, R., Mitchell, K.L., Muller, P., the HRSC Co-Investigator Team, 2005.
Discovery of a flank caldera and very young glacial activity at Hecates Tholus,
Mars. Nature 434, 356–361. doi:10.1038/nature03423.
Hauber, E., van Gasselt, S., Chapman, M.G., Neukum, G., 2008. Geomorphic evidence
for former lobate debris aprons at low latitudes on Mars: Indicators of the
martian paleoclimate. J. Geophys. Res. 113, E02007. doi:10.1029/2007JE002897.
Head, J.W., Marchant, D.R., 2003. Cold-based mountain glaciers on Mars: Western
Arsia Mons. Geology 31, 641–644.
Head, J.W., Marchant, D.R., 2006. Evidence for global-scale northern mid-latitude
glaciation in the Amazonian period of Mars: Debris-covered glacier and valley
glacier deposits in the 30°–50°N latitude band. Lunar Planet. Sci. 37. Abstract
1127.
Head, J.W., Marchant, D.R., 2009. Inventory of ice-related deposits on Mars:
Evidence for burial and long-term sequestration of ice in non-polar regions and
implications for the water budget and climate evolution. Lunar Planet. Sci. 40.
Abstract 1356.
Head, J.W., Pratt, S., 2001. Extensive Hesperian-aged south polar ice sheet on Mars:
Evidence for massive melting and retreat, and lateral flow and ponding of
meltwater. J. Geophys. Res. 106, 12275–12299. doi:10.1029/2000JE001359.
Head, J.W., Wilson, L., 2007. Heat transfer in volcano–ice interactions on Mars:
Synthesis of environments and implications for processes and landforms. Ann.
Glaciol. 45, 1–13.
Head, J.W.and 13 colleagues, 2005. Tropical to mid-latitude snow and ice
accumulation, flow and glaciation on Mars. Nature 434, 346–351.
Head, J.W., Marchant, D.R., Agnew, M.C., Fassett, C.I., Kreslavsky, M.A., 2006.
Extensive valley glacier deposits in the northern mid-latitudes of Mars:
Evidence for Late Amazonian obliquity-driven climate change. Earth Planet.
Sci. Lett. 241, 663–671.
Head, J.W., Marchant, D.R., Kreslavsky, M.A., 2008. Formation of gullies on Mars:
Link to recent climate history and insolation microenvironments implicate
surface water flow origin. Proc. Natl. Acad. Sci. 105, 13258–13263. doi:10.1073/
pnas.0803760105.
Head, J.W., Marchant, D.R., Dickson, J.L., Kress, A.M., Baker, D.M., 2009. Northern
mid-latitude glaciation in the Late Amazonian period of Mars: Criteria for the
C.I. Fassett et al. / Icarus 208 (2010) 86–100
recognition of debris-covered glacier and valley glacier landsystem deposits.
Earth Planet. Sci. Lett., in press. doi:10.1016/j.epsl.2009.06.041.
Hecht, M.H., 2002. Metastability of liquid water on Mars. Icarus 156, 373–
386.
Helbert, J., Reiss, D., Hauber, E., Benkhoff, J., 2005. Limits on the burial depth of
glacial ice deposits on the flanks of Hecates Tholus, Mars. Geophys. Res. Lett. 32,
L17201. doi:10.1029/2005GL023712.
Hiesinger, H., Head, J.W., 2002. Topography and morphology of the Argyre basin,
Mars: Implications for its geologic and hydrologic history. Planet. Space Sci. 50,
939–981.
Holt, J.W.and 11 colleagues, 2008. Radar sounding evidence for buried glaciers in
the southern mid-latitudes of Mars. Science 322, 1235–1238. doi:10.1126/
science.1164246.
Ingersoll, A.P., 1970. Mars: Occurrence of liquid water. Science 168, 972–
973.
Ivanov, B.A., 2001. Mars/Moon cratering rate ratio estimates. Space Sci. Rev. 96, 87–
104.
Kadish, S.J., Head, J.W., Parsons, R.L., Marchant, D.R., 2008. The Ascraeus Mons fanshaped deposit: Volcano–ice interactions and the climatic implications of coldbased tropical mountain glaciation. Icarus 197, 84–109.
Kargel, J.S., 2004. Mars: A Warmer Wetter Planet. Springer Press, London. 557 pp.
Kargel, J.S., Strom, R.G., 1992. Ancient glaciation on Mars. Geology 20, 3–7.
Kerber, L., Head, J.W., 2009. The age of the Medusae Fossae Formation:
Reassessment using lava flow cast and mold contacts. Lunar Planet. Sci. 40.
Abstract 2235.
Kleman, J., 1994. Preservation of landforms under ice sheet and ice caps.
Geomorphology 9, 19–32.
Kleman, J., Borgstrom, I., 1994. Glacial landforms indicative of a partly frozen bed. J.
Glaciol. 40, 255–264.
Knauth, L.P., Burt, D.M., 2002. Eutectic brines on Mars: Origin and possible
relation to young seepage features. Icarus 158, 267–271. doi:10.1006/icar.2002.
6866.
Knighton, A.D., 1981. Channel form and flow characteristics of supraglacial streams,
Austre Okstindbreen, Norway. Arctic Alpine Res. 13, 295–306.
Kowalewski, D.E., Marchant, D.R., Levy, J.S., Head, J.W., 2006. Quantifying low rates
of summer-time sublimation for buried glacier ice in Beacon Valley, Antarctica.
Antarctic Sci. 18, 421–428.
Kreslavsky, M.A., Head, J.W., 2009. Slope streaks on Mars: A new ‘‘wet” mechanism.
Icarus 201, 517–527.
Laskar, J., Correia, A.C.M., Gastineau, M., Joutel, F., Levrard, B., Robutel, P., 2004. Long
term evolution and chaotic diffusion of the insolation quantities of Mars. Icarus
170, 343–364.
Lee, P., Cockell, C.S., Marinova, M.M., McKay, C.P., Rice, J.W., 2001. Snow and ice
melt flow features on Devon Island, Nunavut, Arctic Canada as possible
analogs for recent slope flow features on Mars. Lunar Planet. Sci. 33. Abstract
1809.
Levy, J.S., Head, J.W., Marchant, D.R., 2007. Lineated valley fill and lobate debris
apron stratigraphy in Nilosyrtis Mensae, Mars: Evidence for phases of glacial
modification of the dichotomy boundary. J. Geophys. Res. 112, E08004.
doi:10.1029/2006JE002852.
Levy, J.S., Head, J.W., Marchant, D.R., 2008. The role of thermal contraction crack
polygons in cold-desert fluvial systems. Antarctic Sci. 20, 565–579. doi:10.1017/
S0954102008001375.
Levy, J.S., Head, J.W., Marchant, D.R., 2009. Concentric crater fill in Utopia Planitia:
History and interaction between glacial ‘‘brain terrain” and periglacial mantle
processes. Icarus 202.
Lewis, K.J., Fountain, A.G., Dana, G.L., 1998. Surface energy balance and meltwater
production for a Dry valley glacier, Taylor Valley, Antarctica. Ann. Glaciol. 27,
603–609.
Lewis, K.J., Fountain, A.G., Dana, G.L., 1999. How important is terminus cliff melt?: A
study of the Canada Glacier terminus, Taylor Valley, Antarctica. Global Planet.
Change 22, 105–115.
Li, H., Robinson, M.S., Jurdy, D.M., 2005. Origin of martian northern hemisphere
mid-latitude lobate debris aprons. Icarus 176, 382–394.
Lobitz, B., Wood, B.L., Averner, M.M., McKay, C.P., 2001. Special feature: Use of
spacecraft data to derive regions on Mars where liquid water would be stable.
Proc. Natl. Acad. Sci. 98, 2132–2137.
Lucchitta, B.K., 1981. Mars and Earth: Comparison of cold-climate features. Icarus
45, 264–303.
Madeleine, J.-B., Forget, F., Head, J.W., Levrard, B., Montmessin, F., Millour, E., 2009.
Amazonian northern mid-latitude glaciation on Mars: A proposed climate
scenario. Icarus 203, 390–405.
Malin, M.C., Edgett, K.S., 2000. Evidence for recent groundwater seepage and surface
runoff on Mars. Science 288, 2330–2335.
Malin, M.C.and 13 colleagues, 2007. Context Camera investigation on board the
Mars Reconnaissance Orbiter. J. Geophys. Res. 112, E05S04. doi:10.1029/
2006JE002808.
Mangold, N., 2003. Geomorphic analysis of lobate debris aprons on Mars at Mars
Orbiter Camera scale: Evidence for ice sublimation initiated by fractures. J.
Geophys. Res. 108, 8021. doi:10.1029/2002JE001885.
Mangold, N., Quantin, C., Ansan, V., Delacourt, C., Allemand, P., 2004. Evidence for
precipitation on Mars from dendritic valleys in the Valles Marineris area.
Science 305, 78–81. doi:10.1126/science.1097549.
Mangold, N., Ansan, V., Masson, Ph., Quantin, C., Neukum, G., 2008. Geomorphic
study of fluvial landforms on the northern Valles Marineris plateau, Mars. J.
Geophys. Res. 113, E08009. doi:10.1029/2007JE002985.
99
Marchant, D.R., Head, J.W., 2007. Antarctic dry valleys: Microclimate zonation,
variable geomorphic processes, and implications for assessing climate change
on Mars. Icarus 192, 187–222.
Marchant, D.R., Head, J.W., 2008. Kilometer-thick ice-sheets in the northern midlatitudes in the Amazonian: Analogs from the East Antarctic ice sheet and the
dry valleys. Lunar Planet. Sci. 39. Abstract 2097.
Mars Channel Working Group, 1983. Channels and valleys on Mars. Geol. Soc. Am.
Bull. 94, 1035–1054.
McEwen, A.S.and 14 colleagues, 2007a. Mars Reconnaissance Orbiter’s High
Resolution Imaging Science Experiment (HiRISE). J. Geophys. Res. 112,
E05S02. doi:10.1029/2005JE002605.
McEwen, A.S.and 32 colleagues, 2007b. A closer look at water-related geologic
activity on Mars. Science 317, 1706–1709. doi:10.1126/science.1143987.
Mellon, M.T., Jakosky, B.M., 1995. The distribution and behavior of martian ground
ice during past and present epochs. J. Geophys. Res. 100, 11781–11799.
Milkovich, S.M., Head, J.W., Marchant, D.R., 2006. Debris-covered piedmont glaciers
along the northwest flank of the Olympus Mons scarp: Evidence for lowlatitude ice accumulation during the Late Amazonian of Mars. Icarus 181, 388–
407.
Milliken, R.E., Mustard, J.F., Goldsby, D.L., 2003. Viscous flow features on the surface
of Mars: Observations from high-resolution Mars Orbiter Camera (MOC)
images. J. Geophys. Res. 108, 5057. doi:10.1029/2002JE002005.
Morgan, G.A., Head, J.W., 2009. Sinton crater, Mars: Evidence for impact into a
plateau icefield and melting to produce valley networks at the Hesperian–
Amazonian boundary. Icarus 202, 39–59.
Mouginis-Mark, P.J., 1987. Water or ice in the martian regolith?: Clues from
rampart craters seen at very high resolution. Icarus 71, 268–286. doi:10.1016/
0019-1035(87)90152-7.
Mustard, J.F., Cooper, C.D., Rifkin, M.K., 2001. Evidence for recent climate change on
Mars from the identification of youthful near-surface ground ice. Nature 412,
411–414.
Neukum, G., Jaumann, R., and the HRSC Co-Investigator Team, 2004. HRSC: The High
Resolution Stereo Camera of Mars Express. In: Mars Express: The
Scientific Payload. European Space Agency Special Publication, ESA-SP1240,
pp. 17–36.
Ostrach, L.R., Head, J.W., 2007. Amazonian lobate debris aprons (LDA) in the midlatitudes of Mars: Assessing the role of water ice in their formation and
evolution. Lunar Planet. Sci. 38. Abstract 1100.
Pain, C.F., Clarke, J.D.A., Thomas, M., 2007. Inversion of relief on Mars. Icarus 190,
478–491.
Parsons, R.A., Kreslavsky, M.A., Nimmo, F., 2008. Martian gully slope measurements
made using HiRISE stereo pairs. Lunar Planet. Sci. 39. Abstract 2328.
Pierce, T.L., Crown, D.A., 2003. Morphologic and topographic analyses of debris
aprons in the eastern Hellas region, Mars. Icarus 163, 46–65. doi:10.1016/
S0019-1035(03)00046-0.
Pieri, D.C., 1980. Martian valleys: Morphology, distributions, age, and origin. Science
210, 895–897.
Plaut, J.J., Safaeinili, A., Holt, J.W., Phillips, R.J., Head, J.W., Seu, R., Putzig, N.E., Frigeri,
A., 2009. Radar evidence for ice in lobate debris aprons in the mid-northern
latitudes of Mars. Geophys. Res. Lett. 36, L02203. doi:10.1029/2008GL036379.
Quantin, C., Allemand, P., Mangold, N., Dromart, G., Delacourt, C., 2005. Fluvial and
lacustrine activity on layered deposits in Melas Chasma, Valles Marineris, Mars.
J. Geophys. Res. 110, E12S19. doi:10.1029/2005JE002440.
Rieder, R., Economou, T., Wänke, H., Turkevich, A., Crisp, J., Brükner, J., Dreibus, G.,
McSween, H.Y., 1997. The chemical composition of martian soil and rocks
returned by the mobile alpha proton X-ray spectrometer: Preliminary results
from the X-ray mode. Science 278, 1771–1774.
Safaeinili, A., Holt, J., Posiolova, L., Phillips, R., Head, J.W., Seu, R., and the SHARAD
Team, 2009. New RADAR evidence for glaciers in Mars Phlegra Montes region.
Lunar Planet. Sci. 40. Abstract 1988.
Shean, D.E., Head, J.W., Marchant, D.R., 2005. Origin and evolution of a cold-based
tropical mountain glacier on Mars: The Pavonis Mons fan-shaped deposit. J.
Geophys. Res. 110, E05001. doi:10.1029/2004JE002360.
Shean, D.E., Head, J.W., Fastook, J.L., Marchant, D.R., 2007. Recent glaciation at high
elevations on Arsia Mons, Mars: Implications for the formation and evolution of
large tropical mountain glaciers. J. Geophys. Res. 112, E03004. doi:10.1029/
2006JE002761.
Skidmore, M.L., Sharp, M.J., 1999. Drainage system behaviour of a high-Arctic
polythermal glacier. Ann. Glaciol. 28, 209–215.
Smith, D.E. et al., 1999. The global topography of Mars and implications for surface
evolution. Science 284, 1495–1503.
Squyres, S.W., 1979. The distribution of lobate debris aprons and similar flows on
Mars. J. Geophys. Res. 84 (B14), 8087–8096.
Squyres, S.W., Carr, M.H., 1986. Geomorphic evidence for the distribution of ground
ice on Mars. Science 231, 249–252.
Swanger, K.M., Marchant, D.R., Kowalewski, D.E., Head, J.W., 2010. Viscous flow
lobes in central Taylor Valley, Antarctica: Origin as remnant buried glacial ice.
Geomorphology, Submitted for publication.
Weitz, C.M., Milliken, R.E., Grant, J.A., McEwen, A.S., Williams, R.M.E., Bishop, J.L.,
2008. Light-toned strata and inverted channels adjacent to Juventae and Ganges
chasmata, Mars. Geophys. Res. Lett. 35, L19202. doi:10.1029/2008GL035317.
Williams, R.M.E., Edgett, K.S., 2005. Remnants of the courses of fine-scale,
precipitation-fed runoff streams preserved in the martian rock record. Lunar
Planet. Sci. 36. Abstract 1173.
Williams, R.M.E., Malin, M.C., 2008. Sub-kilometer fans in Mojave Crater, Mars.
Icarus 198, 365–383.
100
C.I. Fassett et al. / Icarus 208 (2010) 86–100
Williams, R.M.E., Chidsey Jr., T.C., Eby, D.E., 2007. Exhumed paleochannels in central
Utah: Analogs for raised curvilinear features on Mars. In: Willis, G.C., Hylland, M.D.,
Clark, D.L., Chidsey, T.C., Jr. (Eds.), Central Utah—Diverse Geology of a Dynamic
Landscape, vol. 36. Utah Geological Association Publication, pp. 221–235.
Williams, R.M.E., Irwin, R.P., Zimbleman, J.R., 2009. Evaluation of paleohydrologic
models for terrestrial inverted channels: Implications for application to martian
sinuous ridges. Geomorphology 107, 300–315. doi:10.1016/j.geomorph.2008.
12.015.