Icarus 216 (2011) 23–39
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Icarus
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Evidence for Amazonian northern mid-latitude regional glacial landsystems
on Mars: Glacial flow models using GCM-driven climate results and comparisons
to geological observations
James L. Fastook a, James W. Head b,⇑, Francois Forget c, Jean-Baptiste Madeleine c, David R. Marchant d
a
University of Maine, Orono, ME 04469, USA
Department of Geological Sciences, Brown University, Providence, RI 02912, USA
Laboratoire de Météorologie Dynamique, Institut Pierre Simon Laplace, Université Paris 6, BP 99, 75252 Paris cedex 05, France
d
Department of Earth Sciences, Boston University, Boston MA 02215, USA
b
c
a r t i c l e
i n f o
Article history:
Received 15 December 2010
Revised 21 March 2011
Accepted 14 July 2011
Available online 16 August 2011
Keywords:
Mars, surface
Mars, climate
Mars, polar geology
a b s t r a c t
A fretted valley system on Mars located at the northern mid-latitude dichotomy boundary contains lineated valley fill (LVF) with extensive flow-like features interpreted to be glacial in origin. We have modeled this deposit using glacial flow models linked to atmospheric general circulation models (GCM) for
conditions consistent with the deposition of snow and ice in amounts sufficient to explain the interpreted
glaciation. In the first glacial flow model simulation, sources were modeled in the alcoves only and were
found to be consistent with the alpine valley glaciation interpretation for various environments of flow in
the system. These results supported the interpretation of the observed LVF deposits as resulting from initial ice accumulation in the alcoves, accompanied by debris cover that led to advancing alpine glacial
landsystems to the extent observed today, with preservation of their flow texture and the underlying
ice during downwasting in the waning stages of glaciation. In the second glacial flow model simulation,
the regional accumulation patterns predicted by a GCM linked to simulation of a glacial period were used.
This glacial flow model simulation produced a much wider region of thick ice accumulation, and significant glaciation on the plateaus and in the regional plains surrounding the dichotomy boundary. Deglaciation produced decreasing ice thicknesses, with flow centered on the fretted valleys. As plateaus lost
ice, scarps and cliffs of the valley and dichotomy boundary walls were exposed, providing considerable
potential for the production of a rock debris cover that could preserve the underlying ice and the surface
flow patterns seen today. In this model, the lineated valley fill and lobate debris aprons were the product
of final retreat and downwasting of a much larger, regional glacial landsystem, rather than representing
the maximum extent of an alpine valley glacial landsystem. These results favor the interpretation that
periods of mid-latitude glaciation were characterized by extensive plateau and plains ice cover, rather
than being restricted to alcoves and adjacent valleys, and that the observed lineated valley fill and lobate
debris aprons represent debris-covered residual remnants of a once more extensive glaciation.
Ó 2011 Elsevier Inc. All rights reserved.
1. Introduction
1.1. Background and evidence for climate change on Mars
Exploration of Mars over the decades has shown that there are
significant volumes of water sequestered in various surface deposits at different latitudes, such as the polar layered terrains (e.g.
Thomas et al., 1992; Keiffer and Zent, 1992; Phillips et al., 2008),
in the shallow subsurface at high latitudes (e.g., Boynton et al.,
2002; Feldman et al., 2002), in the deeper cryosphere (e.g., Clifford,
1993), and even in mid-latitude alpine-valley glacial landsystems
⇑ Corresponding author.
E-mail addresses: fastook@maine.edu (J.L. Fastook), james_head@brown.edu
(J.W. Head).
0019-1035/$ - see front matter Ó 2011 Elsevier Inc. All rights reserved.
doi:10.1016/j.icarus.2011.07.018
(e.g., Head et al., 2010). Furthermore, analysis of the geological record of Mars shows a rich array of Amazonian and Hesperian-aged
features and deposits that involved water, sometimes in huge
amounts (e.g. Clifford, 1993; Carr, 1996; Masson et al., 2001).
Noachian-aged surfaces display valley networks that suggest a
‘‘warmer and wetter’’ early Mars (e.g., Craddock and Howard,
2002; Fassett and Head, 2008). Clearly, the climate has changed
during martian history from a time when liquid water flowed
across the surface in valley networks and pluvial activity may have
occurred (e.g., Craddock and Howard, 2002; Fassett and Head,
2008), to the very cold and hyperarid surface conditions observed
today (e.g., Zurek, 1992; Head et al., 2003a; Carr and Head, 2010).
Coincident with this long-term climate change has been a transition in the global hydrological cycle from an early Mars when it
appears that the hydrological cycle was vertically integrated (e.g.,
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J.L. Fastook et al. / Icarus 216 (2011) 23–39
Carr, 1996; Clifford and Parker, 2001; Baker, 2001; Head et al.,
2003b) to one in which the hydrological cycle is horizontally stratified, with a cryosphere separating the groundwater system from
the surface system. The surface system then consists of three major
water reservoirs: (1) surface deposits (currently the polar caps), (2)
the atmosphere, and (3) the regolith, and most of the movement of
water is associated with seasonal and longer-term climate changes
related to spin-axis/orbital variations (e.g., Laskar et al., 2004).
In addition to these long-term global climate trends, there is
abundant evidence for shorter-term climate fluctuations driving
the surface water cycle related to variations in these spin-axis/
orbital parameters, such as obliquity, eccentricity and precession
(e.g., Laskar et al., 2004). The discovery and documentation of geologically recent gullies (e.g., Malin and Edgett, 2000) and further
documentation of Amazonian tropical mountain glaciers (e.g.,
Williams, 1978; Head and Marchant, 2003; Shean et al., 2005,
2007; Milkovich et al., 2006; Kadish et al., 2008) are among the
growing evidence suggesting that variations in orbital parameters
may cause climate excursions resulting in the redistribution of
water and a Mars very different from that typical of very recent
history. Indeed, Laskar et al. (2004) predict that throughout much
of its history Mars may have been characterized by an obliquity
significantly higher than its current value.
One of these high obliquity states modeled by Forget et al.
(2006) (for obliquity at 45° with a source of moisture at the present
poles) produced regions of persistent ice accumulation on the
northwestern flanks of the Tharsis volcanoes consistent with geological observations (e.g., Head and Marchant, 2003), and glacial
flow models of the resulting ice sheets agreed well with details
of the deposits (Fastook et al., 2008a). Furthermore, geologic analysis of the mid-latitudes of Mars has shown abundant evidence for
Amazonian ice-assisted flow of debris (e.g., Squyres, 1978, 1979;
Lucchitta, 1981, 1984; Pierce and Crown, 2003; Chuang and Crown,
2005; Li et al., 2005).
Fretted terrain and fretted channels, well developed in the
northern part of Arabia Terra (Sharp, 1973), were clearly identified
in Viking data as the location of lobate debris aprons (LDA) and lineated valley fill (LVF) (Squyres, 1978, 1979). Carr and Schaber
(1977) analyzed Viking Orbiter images, concluding that frost creep
and gelifluction were the primary processes in LDA formation.
Squyres (1978) documented the fact that debris aprons (LDA),
characterized by sharply-defined flow fronts and convex-upward
surfaces, extend from massifs and escarpments outward to distances of up to 20 km. Squyres (1978, 1979) interpreted the
aprons to have resulted from mass-wasted debris that flowed
due to the presence of interstitial ice, envisioning that during periods of climate change, atmospheric water vapor would diffuse into
talus piles formed at the base of steep topography, and interstitial
ice would cause mobilization and flow of the talus to produce the
lobes. Longitudinal ridges and grooves in surface debris seen on the
floors of the fretted valleys (LVF) were interpreted to form by iceassisted debris aprons flowing from valley wall talus slopes and
converging in the middle of the valley floor (Squyres, 1978,
1979). A number of workers noted that some LVF in the Nilosyrtis
region appeared consistent with down-valley flow; following earlier ideas by Kochel and Baker (1981). Kochel and Peake (1984)
interpreted some debris aprons as formed by processes similar to
those of terrestrial ice-cored or ice-cemented rock glaciers and further noted a transition from flow-parallel to flow-transverse ridgeand-furrow topography. Lucchitta (1984) interpreted the LDA and
LVF as flow of debris with interstitial ice and showed evidence of
local down-valley flow of LVF, likening some examples to flow patterns in glacial ice in Antarctica. Carr (1996, pp. 116–120) urged
caution in the general interpretation of a range of landforms as
being due to glaciation. He cited: (1) the fact that the glacial
hypotheses requires major climate change late in Mars history
(e.g., Baker et al., 1991) for which he found little supporting evidence and (2) that the low erosion rates in the Amazonian argues
against any major climate excursions accompanied by precipitation, as envisioned by Baker (2001).
More recently, Pierce and Crown (2003) used new image and
altimetry data and found evidence for a wide range of possible
LDA ice content, interpreting the evidence to be consistent with
multiple models of apron formation (e.g., rock glacier ice-assisted
creep of talus, ice-rich landslides, debris-covered glaciers), with
typical LDA resulting from flow of debris that was enriched in
ground ice. Mangold (2003) showed that ice content in LDA may
exceed pore space, and Li et al. (2005) compared LDA MOLA profiles to simple plastic and viscous power law models for ice–rock
mixtures, concluding that LDAs were ice-rich rock mixtures with
some perhaps >40% ice by volume. Chuang and Crown (2005), documenting the detailed character of 65 LDA, concluded that they
consisted of mixtures of debris and ice, but that it was ‘‘difficult
to constrain the internal distribution of ice or the method of debris
apron initiation from the current datasets.’’ In summary, all workers agree that formation of LDA and LVF involve both talus and ice,
but there is disagreement as to the amount of ice: One end-member calls on formation by ice-assisted creep of talus (often defined
as producing a ‘‘rock glacier’’) (e.g., Squyres, 1978, 1979), while another end-member calls on formation as debris-covered glaciers,
(predominantly glacial ice with a cover of sublimation lag or till)
(e.g., Head et al., 2005, 2006a,b).
Head et al. (2010) developed criteria to distinguish between
these end-members and investigated a wide range of occurrences
to assess the origin of LDA and LVF. On the basis of the background
studies described above, and using a range of terrestrial analogs
applicable to the recent cold-desert environment of Mars (e.g.,
Marchant and Head, 2006, 2007), Head et al. (2010) developed
the following criteria to assist in the identification of debris-covered glacial-related terrains on Mars (features, followed by the
interpretation of each, listed in parentheses): (1) alcoves, theatershaped indentations in valley and massif walls (local snow and
ice accumulation zones and sources of rock debris cover); (2) parallel arcuate ridges facing outward from these alcoves and extending down slope as lobe-like features (flow-deformed ridges of
debris); (3) shallow depressions between these ridges and the alcove walls (zones originally rich in snow and ice, which subsequently sublimated, leaving a depression); (4) progressive
tightening and folding of parallel arcuate ridges where abutting
adjacent lobes or topographic obstacles (constrained debris-covered glacial flow); (5) progressive opening and broadening of arcuate ridges where there are no topographic obstacles (unobstructed
flow of debris-covered ice); (6) circular to elongate pits in lobes
(differential sublimation of surface and near-surface ice); (7) larger
tributary valleys containing LVF formed from convergence of flow
from individual alcoves (merging of individual lobes into LVF); (8)
individual LVF tributary valleys converging into larger LVF trunk
valleys (local valley debris-covered glaciers merging into larger
intermontaine glacial systems); (9) sequential deformation of
broad lobes into tighter folds, chevron folds, and finally into lineated valley fill (progressive glacial flow and deformation); (10)
complex folds in LVF where tributaries join trunk systems (differential flow velocities causing folding); (11) horseshoe-like flow
lineations draped around massifs in valleys and that open in a
down-slope direction (differential glacial flow around obstacles);
(12) broadly undulating along-valley topography, including local
valley floor highs where LVF flow is oriented in different down-valley directions (local flow divides where flow is directed away from
individual centers of accumulation); (13) integrated LVF flow systems extending for tens to hundreds of kilometers (intermontaine
glacial systems); (14) rounded valley wall corners where flow
converges downstream, and narrow arete-like plateau remnants
J.L. Fastook et al. / Icarus 216 (2011) 23–39
between LVF valleys (both interpreted to be due to valley glacial
streamlining). Taken together, the occurrence of a number of these
types of features in an area is interpreted to represent the former
presence of debris-covered glaciers and valley glacial systems in
the Deuteronilus–Protonilus region (Head et al., 2010).
Detailed geological analysis of several different areas in the
northern mid-latitudes (Fig. 1a) has shown strong evidence for
the presence of valley glacial landsystems (Eyles, 1983; Evans,
2003) during the Amazonian and the sequestration and preservation of debris-covered glacial ice (Head et al., 2005, 2006a,b,
2010; Levy et al., 2007; Kress and Head, 2008; Morgan et al.,
2009; Dickson et al., 2008, 2010; Hauber et al., 2008; Marchant
and Head, 2008; Baker et al., 2010), interpretations supported by
the detection of extensive buried ice by the SHARAD (SHAllow RADar) instrument on board the Mars Reconnaissance Orbiter (e.g.,
Holt et al., 2008a,b; Plaut et al., 2009, 2010). Further atmospheric
general circulation modeling by Madeleine et al. (2009) for an
obliquity of 35° with the moisture source in the Tharsis region,
rather than the poles, produced persistent ice accumulation along
the mid-latitude region and the dichotomy boundary, also consistent with the presence of these geological features indicative of
valley glaciation documented there.
In this contribution we use the results of the LMD/GCM
(Madeleine et al., 2009) that reproduce conditions that favor ice
accumulation in the mid-latitudes. From these predictions, we
develop glacial flow models to assess the accumulation and flow
patterns of ice to compare to the patterns documented by the
25
geological observations, to test their validity, to further test the
distinctions between the (1) ice-rich debris creep and (2) debris-covered glacier models for LDA and LVF, and if successful, to
learn more about the nature of interpreted Amazonian valley
and regional glaciation.
1.2. Geological observations
A unique fretted terrain, located along the northern midlatitude dichotomy boundary (Fig. 1a), has long been held to
provide clues to the processes responsible for the formation
and degradation of the dichotomy boundary (Sharp, 1973). Parts
of these fretted valleys (Fig. 1b and c) display a multitude of
characteristics typical of integrated valley glacial landsystems
on Earth (e.g., Eyles, 1983; Evans, 2003), including multiple
theater-headed, alcove-like accumulation areas (Fig. 1d); sharp
arete-like ridges typical of glacial erosion; converging patterns
of downslope valley flow (Fig. 1e); valley lineation patterns
typical of folding and shear (Fig. 1e); wrap-around features indicative of flow around obstacles; and broad piedmont-like lobes as
the valley fill extends out into the northern lowlands (Head et al.,
2006a, 2010). These characteristics suggest that the boundary
area was subjected to large-scale glaciation during the Amazonian. Recognition and documentation of this important late-stage
process provides information critical to reconstructing the original dichotomy boundary, and to the understanding of Amazonian
climate history.
Fig. 1. (a) Northern mid-latitudes dichotomy boundary region showing the locations of key published study areas (1: Kress and Head, 2009; 2: Head et al., 2006b; 3: Head and
Marchant, 2009; 4: Morgan et al., 2009; 5: Head et al., 2006a; 6: Baker et al., 2010; 7: Dickson et al., 2008; 8: Levy et al., 2007; 1/7: Head et al., 2010). (b) Fretted channel and
lineated valley fill system in the central Deuteronilus–Protonilus Mensae (DPM) described in Head et al. (2006a) (box 5 in (a)). Viking Orbiter image mosaic and location map.
Boxes and letters in (b) indicate locations of illustrative images in Head et al. (2006a). (c) Map showing the location and trend of small alcoves and lobate flow-like features
(narrow arrows) and broad flow trends of LVF on the valley floor (wide arrows). Compiled using HRSC orbit 1395, THEMIS and MOC images. (d and e) Images illustrating the
nature of alcoves and LVF features associated with the system shown in (b and c). (d) Four adjacent alcoves oriented convex-inward toward the plateau (bottom) with
individual, convex outward lobes extending from each alcove out into the adjacent LVF. Note the convergence of these individual lobate extensions, their compression and
deformation, and their convergence with the general trends of the LVF on the valley floor (left). Location is box A in (b). THEMIS VIS V11208010. (e) Outflow of two ridged
lobes from alcoves (bottom left and right); lobes join major LVF of area C (upper right) near convergence with B (c). Left lobe is swept westward, forming broad arcuate folds;
right lobe is increasingly compressed until it resembles a tight isoclinal fold. Both lobes ultimately merge into the general LVF parallel to the valley walls. Location is box D in
(b). MOC R1702578.
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1.3. Interpretation as a valley glacial landsystem
Detailed analysis of a typical fretted valley network at 34°E,
41°N in the central region between Deuteronilus and Protonilus
Mensae (Fig. 1a, box 5) was undertaken by Head et al. (2006a)
(Fig. 1b and c). This set of fretted valleys extends for over 200 km
in a N–S direction, from deep within the upland plateau to the edge
of the continuous upland scarp at 41–42°N. Using MOLA altimetry, and Viking, THEMIS, HRSC, and MOC images, Head et al.
(2006a) traced the LVF from the narrow parts of the deeper upland
valleys progressively toward the northern lowlands. In the most
inland regions of the fretted valleys analyzed (areas A and B in
Fig. 1b) the valleys are narrow (10–20 km in width), often have
distinctively cuspate (convex outward) shoulder-to-shoulder alcoves along the walls (region A) or larger tributary valleys often
forming hanging valleys along the wall margins (region B). In the
A/B regions (Fig. 1c) the valley floor is about 1500 m above the
northern distal end of the deposit.
In contrast to the sharp parallel valley walls and parallel alongvalley lineated fill typical of some fretted valleys described elsewhere (e.g., Squyres, 1978; Carr, 1995, 2001), these alcoves are
sources of LVF that can be seen to emerge from individual alcoves
as lobes (Fig. 1c and d), descending onto the valley floor, bending
and turning in a downslope direction (Fig. 1c and e), and merging
with adjacent lobes, deforming in the process (Fig. 1c) ultimately
to form the distinctive along-strike ridges of LVF. These types of alcoves, the arete-like borders between them, the arcuate-ridged
lobes extending from them, the progressive compression and folding of the lobes as they converge onto the valley floor and lose their
identity, are all hallmarks of glaciated valley landsystems on Earth
(Eyles, 1983; Benn et al., 2003). In the context of such valley glacial
environments, the alcoves in the DPM area were interpreted to
represent microenvironments where ice and snow accumulate, alcove walls shed debris onto the ice, and debris-covered glaciers
emerge and move downslope into the main valleys, widening the
alcoves and creating aretes over time (Head et al., 2006a). Mapped
trends show the emergence of lobate material from these alcoves
(Fig. 1c, narrow arrows) and the general trends with which they
merge into the larger-scale LVF (wide arrows). The patterns in
areas A, B, and C (Fig. 1c) clearly indicate the convergence typical
of glaciated valley landsystems (Eyles, 1983; Benn et al., 2003)
where abundant small debris-covered glaciers contribute to larger-scale valley glaciers. As the generally north-trending LVF
emerges from areas A to C (Fig. 1c), the individual LVFs merge into
two main trends, areas D and E. LVF C merges with B and continues
95 km to the NNW where it bifurcates around a mesa to form LVF
F and G, each continuing 50–75 km to the northern lowlands. Western portions of LVF B turn west into D, merging with LVF A to produce complex fold patterns. Portions of A turn NW and extend
almost 100 km to the northern lowlands. Several basic trends are
observed over this region extending from A–B–C to the northern
lowlands (Head et al., 2006a). The elevation of the LVF floor decreases by 1500–2000 m from the proximal areas (areas A–C)
to the end of the LVF in the northern lowlands. LVF slopes tilt toward the north and are less than 1° throughout most of the LVF
floor, but increase rapidly to >1–2° in the northern 25–50 km of
the LVF adjacent to the lowlands. These areas are characterized
by piedmont-like lobes extending into the northern lowlands between mesas. These features are reminiscent of the fronts of terrestrial glaciers in terms of their lobate nature and relatively steep
front, and in many cases are characterized by flow lines, pits and
moraine-like ridges.
In summary, spacecraft data show compelling evidence for an
integrated picture of LVF formation (Fig. 1b–e; Head et al.,
2006a) with a significant role being played by regional valley glaciation (e.g., Eyles, 1983; Post and Lachapelle, 2000; Benn et al.,
2003) in the modification of the valley systems. There is evidence
for: (1) localized alcoves, which are the sources of dozens of narrow, lobate concentric-ridged flows interpreted to be remnants of
debris-covered glaciers; (2) depressions in many current alcoves,
suggesting loss of material from relict ice-rich accumulation zones;
(3) narrow pointed plateau ridge remnants between alcoves, similar to glacially eroded aretes; (4) horseshoe-shaped ridges up-valley and upslope of topographic obstacles; (5) convergence and
merging of LVF fabric in the down-valley direction, and deformation, distortion and folding of LVF in the vicinity of convergence,
all consistent with glacial-like, down-valley flow (Fig. 1c); and
(6) distinctive lobe-shaped termini where the LVF emerges into
the northern lowlands, with associated pitting and concentric terminal ridges.
This assemblage of features can be traced and mapped in an
integrated fashion (Fig. 1b and c) for over 200 km in length, and
covers an area (30,000 km3), 10 times that of the Antarctic
Dry Valleys (Marchant and Head, 2007), being more comparable
in scale to the remnant North American ice caps on Ellesmere
and Baffin Islands in the Canadian high Arctic. This assemblage of
features suggests that the LVF in this region formed as part of a
broad integrated valley glacial system extending for about
200 km in the S–N direction. The extensive development of similar
lineated valley fill occurrences, often with very similar characteristics, in fretted valleys in the DPM region (Fig. 1a; see also Head
et al., 2010), suggests that glaciation may have been instrumental
in the modification of the dichotomy boundary over an area as
large as several million km2. The age of the LVF in different parts
of the DPM region has been estimated as Amazonian (e.g., McGill,
2000; Mangold, 2003; Morgan et al., 2009; Baker et al., 2010), with
relatively younger alteration by processes of mantle deposition and
surface layer sublimation and modification (e.g., Mangold, 2003;
Head et al., 2003a,b; Levy et al., 2009). On the basis of these types
of observations, Head et al. (2006a) (and others; Fig. 1a) interpreted the process of alpine-style valley glaciation to have been
extensive in this region.
To assess this glacial origin hypothesis for this valley system
(Fig. 1b) and the region in general (Fig. 1a), Madeleine et al.
(2009) explored the conditions under which GCMs reproduced sufficient ice deposition and seasonal preservation to cause glaciation.
They found that the LMD/GCM for a dusty Mars atmosphere with
obliquity set to 35° and a water source in the Tharsis region could
generate ice accumulation in good agreement with these geological observations. According to Madeleine et al. (2009), this
proposed climate is what one might expect to follow a higherobliquity excursion (Forget et al., 2006; Levrard et al., 2004) of
the sort thought to build ice sheets on the flanks of the Tharsis volcanoes (Head and Marchant, 2003). We now use the predictions of
the LMD/GCM to formulate glacial flow models to compare to the
patterns of LDA and LVF interpreted by Head et al. (2006a) to be
due to glaciation, and to test the validity of these interpretations
of valley glacial landsystems (Fig. 1a), and to test further the
distinctions between the (1) ice-rich debris creep and (2) debriscovered glacier models for LDA and LVF. We then explore the
implications of the locations of ice depocenters predicted by the
LMD/GCM for more regional glacial landsystems.
2. Glacial flow modeling of the Deuteronilus–Protonilus Mensae
(DPM) region
2.1. The University of Maine Ice Sheet Model (UMISM)
UMISM is an adaptation for the martian environment (Fastook
et al., 2004, 2005, 2006) of a terrestrial ice sheet model used for
time-dependent reconstructions of Antarctic, Greenland, and
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J.L. Fastook et al. / Icarus 216 (2011) 23–39
600
Positive Mass Balance Regions
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Fig. 2. DPM lineated valley fill system showing alcove regions of positive mass balance of 1 mm/a; 0.1 mm/a negative mass balance elsewhere. Background is MOLA gridded
topography, color-coded in elevation (m). Red lines show location of positive mass balance regions. Grid is model reference grid in 10 km boxes.
paleo-ice sheet evolution in response to changing climate on Earth
(Fastook, 1993). UMISM uses a thermo-mechanically coupled Shallow-Ice Approximation (vertically-integrated momentum combined with continuity) where the dominant stress is internal
shear and longitudinal stresses are neglected. Primary input to
the model is the bed on which the ice sheet is to be reconstructed
and the net annual surface mass balance (SMB), or accumulation
rate. Secondary input includes the mean-annual surface temperature and the geothermal heat flux, used to calculate internal temperatures from which the mechanical properties of the ice are
obtained. In addition, internal temperatures allow for the possibility that the base of the ice reaches the pressure melting point, at
which point some sliding criteria can be invoked, a phenomena
we have not observed in any modeled martian glaciers (Fastook
et al., 2008a,b).
The fact that with the exception of the bed, these inputs are
poorly constrained for Mars introduces uncertainty into the results
that will be presented. However, choice of reasonable values for
mean annual temperature, geothermal flux, and SMB still allow production of results comparable to the geologic observations. Given
these uncertainties the choice of a Shallow-Ice Approximation model is also appropriate, since the considerable computational load of a
higher-order model would not produce more accurate results.
2.2. Modeling: alcove-only accumulation areas
As a preliminary step in the modeling of the DPM interpreted
glaciation (Head et al., 2006a), we first examine a simple case
where accumulation only occurs in the alcoves along the valley
walls. This simple treatment has terrestrial analogs in the Dry Valleys of Antarctica (Marchant and Head, 2004, 2005, 2006, 2007). In
particular, we follow a model of ice deposition observed in the Dry
Valleys of Antarctica on Mullins Glacier, a debris-covered glacier
that originates in an alcove analogous to the martian alcoves
(Kowalewski et al., 2011). Here deposition of ice occurs only in a
very limited area at the base of a scarp that also serves as the
source of the surface debris. Without this protective layer of rock
debris, the glacier would sublimate rapidly as it flows down into
the Dry Valleys, an area where normal sublimation removes any
ice exposed at the surface. Marchant and Head (2004, 2005,
2006, 2007) and Marchant et al. (2010) have shown that buried
ice in the Mullins Glacier may be several millions of years old, a situation that would be not be possible without the armoring effect of
the debris cover (Kowalewski et al., 2006).
In reconstructions of Tharsis Montes ice sheets, we used either a
parameterization of the SMB as a function of height (Fastook et al.,
2004, 2005) or GCM results (Fastook et al., 2006) as the source of
mass in the continuity equation. Here we chose the expedient of
simply specifying a positive SMB of +1 mm/year in the alcoves,
and a negative SMB 1/10th of that outside the alcoves (Fig. 2). These
values are typical of debris-covered glaciers in the Dry Valleys, Antarctica (Kowalewski et al., 2011). The shape of the resulting profile is
relatively insensitive to the magnitude of the SMB. SMB is raised to
the 1/8th power in analytic solutions for uniform SMB elliptical profiles, implying a thickening of only 10% for a doubling of the SMB.
What larger or smaller SMB will produce is different formation times
and resulting velocity magnitudes. Doubling SMB effectively doubles velocity because thickness is only increased by 10% and the increased flux must be accommodated by increased velocity.
Changing SMB does not significantly affect the orientation of the
flow field because flow direction is down the surface gradient and
surface changes are small relative to changes in the SMB. These
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J.L. Fastook et al. / Icarus 216 (2011) 23–39
Thickness [3006]
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Fig. 3. DPM lineated valley fill system color-coded thicknesses (m) at 300, 500, 1000, and 1500 Ka. Contour lines indicate surface elevation.
higher or lower velocities will result in faster or slower advance of
the glacier during the formation, and of course will affect retreat
times during collapse. There are no constraints on either the formation times or the velocity magnitudes, so the simple expedient of
+1 mm/a in the alcoves and 0.1 mm/a elsewhere was chosen to
yield a 10:1 ratio between accumulation area and ablation area.
Starting with no ice, the model is run for 2 Ma, which delivers a
flow pattern (primarily the orientation of the velocity field) that
can be compared to the Head et al. (2006a) glacial interpretation.
Resulting ice thicknesses and ice flow velocities, with superimposed surface elevations, are shown in Figs. 3 and 4 for four time
intervals during the growth of these ice deposits.
(1) At 300 Ka (Fig. 3, upper left), the ice is thin (less than 300 m;
blue1 areas) and the flow from the sides is at low velocity (less
than 50 mm/a; Fig. 4, upper left). The flow has not yet
merged along the center of the valleys (blue areas in Fig. 3
upper left). A configuration such as this would not yet produce
the turning flow observed by Head et al. (2006a) (Fig. 1).
(2) By 500 Ka (Fig. 3, upper right), the beginning of a coherent
down-valley flow is observed (green and blue areas in the
lower right-hand and central part). The ice flowing from
1
For interpretation of color in Figs. 1–10 and 12, the reader is referred to the web
version of this article.
29
J.L. Fastook et al. / Icarus 216 (2011) 23–39
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Fig. 4. DPM lineated valley fill system color-coded velocities (mm/a) at 300, 500, 1000, and 1500 Ka. Contours indicate surface elevation.
each side of these parts of the valleys has merged in the center. Note that the velocities (Fig. 4, upper right) are not yet
appreciable, with most velocities being less than 50 mm/a.
(3) By 1000 Ka (Fig. 3, lower left), there is a well-established valley glacier system extending from the fretted valleys in the
south to the mouths of the valleys along the dichotomy
boundary, with thicknesses exceeding 400–500 m in the
southern part of the valley system. Velocities (Fig. 4, lower
left) are almost everywhere higher than 50 mm/a, and
locally in excess of 200 mm/year.
(4) As time progresses to 1500 Ka (Fig. 3, lower right), the valley
glaciers become very well developed in the valleys and
extend out of the valleys and onto the northern lowlands.
Ice thicknesses are locally in excess of 600 m. Velocities
(Fig. 4, lower right) are broadly in excess of 100 mm/a, and
locally exceed 200 mm/a.
Comparison of time steps in Fig. 3 shows that ice accumulation
and flow at either 1000 Ka (Fig. 3, lower left) or 1500 Ka (Fig. 3,
lower right) would clearly be producing the kinds of landforms
30
J.L. Fastook et al. / Icarus 216 (2011) 23–39
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Fig. 5. The region furthest from the dichotomy boundary and the valley mouths (area BC in Fig. 1b, THEMIS VIS V11208010) is characterized by eastward-opening adjacent
alcoves with multiple convex-outward concentric-ridged lobes extending, converging, deforming, and merging with the valley floor LVF trends. Here the model at 1000 and
1500 Ka shows extremely low surface slopes (surface contours are only 50 m). Ice is 300–500 m thick, and velocities are generally less than 50 mm/a. The flow is directed
away from the alcoves with little indications of turning down-valley.
observed by Head et al. (2006a) and interpreted to be glacial in origin (compare to patterns in Fig. 1b–e).
We can further examine these correlations and test the glacial
interpretation by comparing smaller areas of accumulation and
flow orientations to the patterns observed by Head et al. (2006a).
High resolution excerpts from the glacial flow model results, for
both thickness and velocity at 1000 Ka and at 1500 Ka (Figs. 3
and 4), are compared to patterns in the THEMIS images of regions
BC, F, and G from Fig. 1b (Head et al., 2006a) in Figs. 5–7. While
direction of flow is not explicitly shown in Figs. 5–7, the velocity
vector is always perpendicular to the surface elevation contours
and aligns well with the interpreted flow feature orientations.
The region furthest from the dichotomy boundary and the valley mouths (area BC in Fig. 1b) shown in Fig. 5 is characterized by
eastward-opening adjacent alcoves with multiple convex-outward
concentric-ridged lobes extending, converging, deforming, and
merging with the valley floor LVF trends. Here the model at 1000
and 1500 Ka shows extremely low surface slopes (surface contours
are only 50 m). Ice is 300–500 m thick, and velocities are generally
less than 50 mm/a. This matches well with the interpretation of
alcove-sourced debris-covered glaciers merging gradually into
the valley flow system. Velocities are very low and there is little
indication of the flow turning down-valley.
A region further down the valley toward the dichotomy
boundary (area F in Fig. 1b) containing an eastward-facing zone
of multiple alcoves and converging lobate flows in the distal regions of the system, along the edge of the northernmost large
mesa (area G in Fig. 1b) is shown in Fig. 6. Here concentric-outward ridges extend from alcoves and are progressively compressed, folded and flattened as the ridges deform and become
part of the valley floor LVF. In this case, model thicknesses at
1000 and 1500 Ka are 300–500 m and velocities are locally
250 mm/a toward the center of the valley, but less than
100 mm/a in the alcove and adjacent area. In this example,
one can clearly see the low flow rates at the alcove and the
abruptly increasing flow rates at the point of the turning as
the glacial lobes merge with the faster along-valley flow.
A region further down the valley near the dichotomy boundary (area G in Fig. 1b) is away from the alcoves and in the central part of the valley where ice flow was interpreted by Head
et al. (2006a) to have been down-valley glacial flow that bifurcated around massifs in the central part of the valley system
(Fig. 7). LVF bifurcates (bottom right) and flows around a massif
forming a broad up-flow collar and a diffuse, down-flow ‘‘wake’’.
LVF in the narrow pass between massifs is compressed. The lobate LVF extends into the northern lowlands at top left. Note
the arcuate, piedmont-like configuration of the lobe and the high
density of pits. The modeled flow clearly narrows as it pinches
between the two bedrock highs and accelerates in these areas
to values in excess of 200 mm/a.
In summary, the glacial flow and velocity orientation patterns
in the ice flow model simulation clearly match many of the features observed in the lineated valley fill, interpreted by Head
et al. (2006a) to be of glacial origin. Particularly significant were:
(1) the turning flow observed in the lower valley emanating from
the alcoves and merging with the general down-valley flow
(Fig. 6); (2) the deflection of flow around obstacles (consistent with
a relatively thin glacial ice deposit confined to the valley) (Fig. 7);
thicker, more extensive ice sheets might follow the terrain less
faithfully, as the obstacles would be overridden; and (3) the similarities in the thicknesses and extents of the modeled ice and the
LVF deposits (Figs. 1b, c and 3).
We conclude that the glacial flow models described here
successfully reproduce the lineated valley fill patterns and flow
31
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Fig. 6. A region further down the valley toward the dichotomy boundary (area F in Fig. 1b, THEMIS VIS V11208010 with model thickness and velocity at 1000 and 1500 Ka)
containing an eastward-facing zone of multiple alcoves and converging lobate flows in the distal regions of the system, along the edge of the northernmost large mesa (area G
in Fig. 1b). Here concentric-outward ridges extend from alcoves and are progressively compressed, folded and flattened as the ridges deform and become part of the valley
floor LVF. In this case, thicknesses are 300–500 m and velocities are locally 250 mm/a toward the center of the valley, but less than 100 mm/a in the alcove and adjacent
area. In this example, one can clearly see the low flow rates at the alcove and the increased flow rates at the point of the turning of the glacial lobes and the merging with the
along-valley flow.
scenarios interpreted by Head et al. (2006a) to be evidence for alpine-type valley glacial landsystems, and thus support this interpretation by providing a physically-reasonable scenario that
produces results quantitatively similar to the geological observations. In this scenario, ice accumulates in alcoves and other protected areas, and debris cover derived from the adjacent exposed
cliffs records the flow features of the underlying ice even after
the ice beneath has partly sublimated away (Head et al., 2006a).
A conclusion is that we would only expect to see such debris features where a source of surface debris was available, as is observed
with the Mullins Glacier in the Dry Valleys of Antarctica (Marchant
and Head, 2007; Marchant et al., 2010; Kowalewski et al., 2011;
Shean and Marchant, 2010) and in the valley walls of the fretted
terrain and along the dichotomy boundary. Otherwise, if not covered by debris, the ice at these latitudes would have sublimated
away in the modern climate (e.g., Hauber et al., 2008) leaving no
trace.
A corollary to this is that we would predict that any ice deposited at higher levels on the plateau above the valleys, with no higher scarps or nunataks from which debris could be deposited, would
leave no record of ice flow. This raises the question: Could the currently observed LVF (Fig. 1b–e) represent not the greatest lateral
extent of an alpine-type valley landsystem, but instead represent
the waning stages and a remnant of a larger ice sheet preserved
only when the scarps were exposed during retreat (Marchant and
Head, 2006, 2007, 2008)? In the following section we examine the
specific accumulation areas predicted by the LMD/GCM to assess
whether the LVF might be the record of the waning stages of glaciation, after the regional ice surface in the valleys had dropped below the levels of the surrounding scarps, allowing debris to
accumulate on the surface of the glaciers.
3. Glacial flow modeling of GCM-defined accumulation areas
3.1. The Mars GCM of the Laboratoire de Météorologie Dynamique
In order to specify the spatial distribution of the SMB of accumulated ice for a glacial flow model, one can arbitrarily choose values, and explore the consequences for specific values and
predictions as we have successfully done above. An alternative approach is to use the results of a GCM that was chosen on the basis
of its ability to reproduce the broad conditions that are observed
geologically, as we did in the Tharsis region (e.g., Head and
Marchant, 2003; Forget et al., 2006; Fastook et al., 2004, 2005,
2008a). The GCM used is the LMD/GCM (Laboratoire de Météorologie Dynamique, Forget et al., 1999); this GCM is able to reproduce
the present-day water cycle with good accuracy (Montmessin
et al., 2004). Extensive exploration of the parameter space found
that necessary conditions for persistent ice deposition along the
northern mid-latitude dichotomy boundary (Fig. 8a) included
moderate obliquity (25–35°), high eccentricity (0.1) with perihelion at Lp = 270°, high dust opacity (1.5–2.5), and a water source
from sublimation of an ice sheet deposited on the flanks of the
Tharsis volcanoes during a prior period of higher obliquity where
tropical mountain glaciers are observed in the geological record
(Head and Marchant, 2003; Shean et al., 2005, 2007; Milkovich
et al., 2006; Kadish et al., 2008). The high dust content of the atmosphere was necessary to increase its water vapor holding capacity,
thereby moving the saturation region to the northern mid-latitudes. Precipitation events are then controlled by topographic forcing of stationary planetary waves and transient weather systems,
producing surface ice distribution and amounts that are consistent
with the geological record. Both lower eccentricity and reversed
32
J.L. Fastook et al. / Icarus 216 (2011) 23–39
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Fig. 7. A region further north near the dichotomy boundary (area G in Fig. 1b, THEMIS VIS V11208010 with model thickness and velocity at 1000 and 1500 Ka) far from the
alcoves and in the central part of the valley where ice flow is down-valley. LVF bifurcates (bottom right) and flows around a massif forming a broad up-flow collar and a
diffuse, down-flow ‘‘wake’’. LVF in the narrow pass between massifs is compressed. The lobate LVF extends into the northern lowlands at top left. Note the arcuate, piedmontlike configuration of the lobe and the high density of pits. Model results show the flow clearly narrowing between the two bedrock highs and increasing the velocity to values
in excess of 200 mm/a.
perihelion did not produce persistent ice cover consistent with
geological observations. This use of a GCM tuned to produce a positive SMB where glaciers are interpreted to have existed, informs
us about the necessary state of the climate at the time the glaciers
were deposited.
3.2. Specific global circulation model results for the northern midlatitude dichotomy boundary area
Work with the LMD/GCM (Forget et al., 1999; Montmessin et al.,
2004; Levrard et al., 2004; Hourdin et al., 1993) has provided a
framework so that a map of potential accumulation rates for the
dichotomy boundary region can be produced (Madeleine et al.,
2009), which can then be used in an ice sheet model to describe a
possible ice sheet with associated valley glaciation observed in the
glacial geology (Head et al., 2006a; Fastook et al., 2008b). The distribution of positive SMB regions for the Madeleine et al. (2009) climate simulation showing best conditions for development of the
mid-latitude glaciation are illustrated in Fig. 8a. The valley glaciation
region described and tested above (Head et al., 2006a) and further
discussed here lies in the area designated by the arrow in Fig. 8a
(see also Fig. 1a, box 5). With peak values of ice accumulation reaching 16 mm/a, this pattern of SMB is clearly capable of producing a
large ice sheet along the dichotomy boundary, but will its behavior
agree with the geological observations (Fig. 1)?
3.3. Ice sheet modeling results
Clearly such a wide area of positive SMB will create a broad,
extensive ice sheet (Marchant and Head, 2008), as opposed to only
the localized valley glaciers observed in the geological record
(Fig. 1a; Head et al., 2006a,b, 2010; Levy et al., 2007; Kress and
Head, 2009; Morgan et al., 2009; Dickson et al., 2008, 2010; Hauber
et al., 2008; Baker et al., 2010). We chose to focus attention again
on the Deuteronilus–Protonilus Mensae (DPM) valley system described in Head et al. (2006a) (Fig. 1b–e). Running the ice sheet
model at a resolution sufficient to resolve the valley, however, is
prohibitive for such a broad area. We utilize instead the embedded-grid feature of UMISM. This feature allows us to run a
broad-domain, low-resolution grid with a more limited-domain,
higher-resolution grid embedded within it. This embedded grid obtains boundary condition information from a spatial and temporal
interpolation of the low-resolution grid. This embedded feature allows us to nest grids, so that the jump in resolution is not so extreme as to produced spurious results.
The grids in the nest used in this model, with topography from
MOLA, are also shown in Fig. 8. Starting with Fig. 8b, the broad grid
outlined by the box in Fig. 8a (with 10,164 nodes and a resolution
of 50 km), the nest progresses to Fig. 8c (with 16,625 nodes and
12 km resolution), onto Fig. 8d (with 7521 nodes and 6 km resolution), and finally to Fig. 8e, with the highest resolution (1.6 km, and
12,769 nodes). The three outer grids in the nest at 600 Ka are
shown in Fig. 9, the time at which growth is stopped and retreat
begins. The three vertical columns contain surface elevation, ice
thickness, and ice velocity, respectively. The three rows correspond
to the three outermost grids (Fig. 8b–d).
On the basis of this glacial accumulation and flow model, it is
observed that the broad pattern of positive SMB in the northern
plains (Fig. 8a) builds an ice sheet up to 4 km thick with a volume
close to 9 million km3 in 600 Ka (Fig. 9, first row, left and middle
33
J.L. Fastook et al. / Icarus 216 (2011) 23–39
(a)
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Fig. 8. (a) Mass balance in mm/yr from Madeleine et al. (2009) showing persistent ice deposition along the northern mid-latitude dichotomy boundary. GCM parameters
included moderate obliquity (25–35°), high eccentricity (0.1) with perihelion at Lp = 270°, high dust opacity (1.5–2.5), and a water source from sublimation of an ice sheet
deposited on the flanks of the Tharsis volcanoes; (b) the outermost of the nested grids with a resolution of 50 km, outlined by the box in (a); (c) with 12 km resolution,
outlined in (b); (d) with 6 km resolution, outlined in (c); and (e) with the highest resolution of 1.6 km, outlined in (d).
columns). This is considerably more than our own estimate of a
Tharsis-region volume of 0.53 million km3 (Fastook et al.,
2008a,b), but the Tharsis-region estimate was very conservative
and was constrained to be no greater than the clearly observable
glacial deposits mapped in that area. Other estimates of the Tharsis
volumes are considerably higher with much more extensive ice
sheets. One reconstruction by Kite and Hindmarsh (2007) is 25–
54 million km3, a size hard to reconcile with the current volume
of the North Polar Layered Deposits, which is 1.2–1.7 million km3
(Zuber et al., 1998). If we restrict our ice sheet to the next grid in
the nest (Figs. 8c and 9, second row) our volume is close to 2 million km3, still larger than our estimate for Tharsis, but closer to the
volume of the current North Polar Layered Deposits. We do in
fact take depletion of the source into account, as that is the reason we turn off the accumulation portion of the SMB at 600 Ka
when retreat and collapse of the ice sheet begins, leading to
the patterns of ice thickness shown in the valley complex in
Fig. 12. Clearly, the actual sources, volumes and transport pathways of water ice are not fully understood and subject to current
investigation (e.g., Mischna et al., 2003; Levrard et al., 2004,
2007). Nonetheless, the broad modeled ice sheet covers much
of the adjacent plateau and is particularly thick near the dichotomy boundary (Fig. 9, middle column); fretted valleys, particularly in areas interpreted to be the sites of local valley glacial
34
J.L. Fastook et al. / Icarus 216 (2011) 23–39
SURFACE [1500] {a}
THICKNESS [1500] {a}
TIME= 600000
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-5
-10
-4
0
VELOCITY [1500] {a}
TIME= 600000
10
-3
-2
20
30
-1
0
40
50
1
2
60
3 0.0
0.5
-10
70
TIME= 600000
1.0
1.5
0
10
2.0
2.5
20
30
3.0
3.5
40
4.0
50
4.5
5.0
60
5.5 -10
-9
-10
70
-8
0
-7
10
-6
-5
-4
20
30
-3
-2
-1
40
50
0
1
2
60
3
70
60
60
60
60
50
50
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50
40
40
40
40
30
30
30
30
-10
0
10
20
30
40
50
60
70
-10
SURFACE [2100] {b}
-5
-4
20
0
10
20
30
40
50
60
-10
70
THICKNESS [2100] {b}
TIME= 600000
-6
20
20
20
20
-2
30
-1
0
1
40
2
3 0.0
0.5
50
1.0
20
10
20
30
40
50
60
70
VELOCITY [2100] {b}
TIME= 600000
-3
0
TIME= 600000
1.5
2.0
2.5
30
3.0
3.5
40
4.0
4.5 5.0
50
5.5 -10
-9
-8
-7
-6
20
-5 -4
30
-3
-2 -1
40
0
1
50
2
3
50
50
50
50
40
40
40
40
30
30
20
30
40
30
50
20
SURFACE [3100] {c}
-5
-4
30
40
30
50
20
THICKNESS [3100] {c}
TIME= 600000
-6
30
-2
-1
35
0
1
2
3 0.0
40
0.5
1.0
40
50
VELOCITY [3100] {c}
TIME= 600000
-3
30
TIME= 600000
1.5
2.0
2.5
30
3.0
3.5
35
4.0
4.5
5.0
5.5-10
40
-9
-8
-7
30
-6
-5
-4
-3
-2
-1
0
1
35
2
3
40
45
45
45
45
40
40
40
40
30
35
40
30
35
40
30
35
40
Fig. 9. Surface elevation, thickness, and velocity (columns) for the three outermost grids (rows) of Fig. 8b–d in the DPM lineated valley fill system.
landsystem deposits (e.g., Fig. 1a; Morgan et al., 2009; Head
et al., 2006b; Baker et al., 2010; Dickson et al., 2008; Levy
et al., 2007), show thick ice accumulations.
The specific DPM valley (Head et al., 2006a; Fig. 1b–e) lies on a
GCM grid point where high accumulation rates are predicted (the
resolution of the GCM is 5.625° longitude by 3.75° latitude). This
results in a ‘‘peninsula’’ of ice that stretches into the highlands of
the dichotomy boundary with thicknesses approaching 3 km. Note
that in the logarithmic velocity scale, a value of 1 corresponds to a
velocity of 10 mm/a. The DPM valley (Fig. 1b–e) lies in a saddle region where ice flow is clearly faster, 100 mm/a, and is channelized in the trunk of the valley.
This channelization is more evident in the highest-resolution
grid (Fig. 8e). In Fig. 10 the evolution during growth of the valley
ice complex (top to bottom: surface, thickness, and velocity) is
shown at 100 Ka intervals (left to right: 100–600 Ka). In the surface
35
J.L. Fastook et al. / Icarus 216 (2011) 23–39
SURFACE [4001] {d}
SURFACE [4002] {d}
TIME= 100000
-6 -5
32
-4
-3
33
-2
SURFACE [4003] {d}
TIME= 200000
-1 0
34
1
2
35
3 -6 -5
32
-4
-3
33
-2
SURFACE [4004] {d}
TIME= 300000
-1 0
34
1
2
35
3 -6 -5
32
-4
-3
33
-2
-1 0
34
1
2
35
3 -6 -5
32
-4
-3
33
-2
SURFACE [4006] {d}
SURFACE [4005] {d}
TIME= 400000
TIME= 600000
TIME= 500000
-1 0
34
1
2
35
3 -6 -5
32
-4
-3
33
-2
-1 0
34
1
2
35
3 -6 -5
32
-4
-3
33
-2
-1 0
34
1
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3
43
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40
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THICKNESS [4001] {d}
33
34
35
THICKNESS [4002] {d}
TIME= 100000
40
32
33
34
35
THICKNESS [4003] {d}
TIME= 200000
40
40
32
33
34
35
32
THICKNESS [4004] {d}
TIME= 300000
33
34
THICKNESS [4005] {d}
TIME= 400000
40
32
35
33
34
35
THICKNESS [4006] {d}
TIME= 500000
TIME= 600000
0.0 0.5 1.0 1.5 2.0 2.5 3.0 3.5 4.0 4.5 5.0 5.5 0.0 0.5 1.0 1.5 2.0 2.5 3.0 3.5 4.0 4.5 5.0 5.5 0.0 0.5 1.0 1.5 2.0 2.5 3.0 3.5 4.0 4.5 5.0 5.5 0.0 0.5 1.0 1.5 2.0 2.5 3.0 3.5 4.0 4.5 5.0 5.5 0.0 0.5 1.0 1.5 2.0 2.5 3.0 3.5 4.0 4.5 5.0 5.5 0.0 0.5 1.0 1.5 2.0 2.5 3.0 3.5 4.0 4.5 5.0 5.5
32
33
34
35
32
33
34
35
32
33
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35
33
34
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32
33
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35
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33
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40
40
32
33
34
40
32
35
VELOCITY [4001] {d}
33
34
35
VELOCITY [4002] {d}
TIME= 100000
40
32
33
34
35
VELOCITY [4003] {d}
TIME= 200000
40
32
33
34
35
VELOCITY [4004] {d}
TIME= 300000
40
40
32
33
34
35
32
VELOCITY [4005] {d}
TIME= 400000
33
34
35
VELOCITY [4006] {d}
TIME= 500000
TIME= 600000
-10 -9 -8 -7 -6 -5 -4 -3 -2 -1 0 1 2 3 -10 -9 -8 -7 -6 -5 -4 -3 -2 -1 0 1 2 3 -10 -9 -8 -7 -6 -5 -4 -3 -2 -1 0 1 2 3 -10 -9 -8 -7 -6 -5 -4 -3 -2 -1 0 1 2 3 -10 -9 -8 -7 -6 -5 -4 -3 -2 -1 0 1 2 3 -10 -9 -8 -7 -6 -5 -4 -3 -2 -1 0 1 2 3
32
33
34
35
32
32
33
34
35
33
35
32
34
33
35
34
35
32
33
34
35
32
33
34
43
43
43
43
43
43
43
42
42
42
42
42
42
42
41
41
41
41
41
41
41
40
33
34
35
40
40
40
32
32
33
34
35
32
33
34
35
40
32
33
34
35
40
40
32
33
34
35
32
33
34
35
Fig. 10. Growth to 600 Ka, showing surface (m), thickness (m), and velocity (log 10(mm/a)) at 100, 200, 300, 400, 500, and 600 Ka in the DPM lineated valley fill system.
figures (Fig. 10, top row), we see a gradual progressive drowning of
the valley topography. Thickness (Fig. 10, middle row) begins with
a uniform mantling, but quickly evolves into thicker ice in the valleys (4 km) but much thinner over the plateaus (less than 3 km)
and aretes (1 km). Recalling that ice flow directions are down
the surface topographic gradient, perpendicular to surface elevation contours, we see the organization of a coherent flow pattern
with maximum velocity (Fig. 10, lower row) in the valleys
(100 mm/a) but one, two, and even three orders of magnitude
less over the more thinly covered aretes and plateaus between
the valley trunks. Thus, to a first order, the general distribution
of ice accumulation, as predicted by the LMD/GCM, readily
reproduces regional plateau glaciation, and shows that glacial flow
is focused into the fretted terrain valleys due to the effect of the
pre-existing topography.
What is the fate of the regional ice sheet during its sublimation
and retreat? As was mentioned, with the relatively high accumulation rates predicted by Madeleine et al. (2009), growth of a significant ice sheet is relatively rapid, and we grew the ice sheet under
steady climate conditions for 600 Ka. At this point, we removed the
positive accumulation component of the SMB, leaving only the
negative sublimation component. This might be expected to occur,
Fig. 11. Ice evolution versus time for a typical point in Deuteronilus (45°N28.1°E).
By tracking increasing and decreasing thickness, the accumulation and ablation
components of the mass balance can be separated.
36
J.L. Fastook et al. / Icarus 216 (2011) 23–39
SURFACE [4102] {d}
SURFACE [4103] {d}
TIME= 800000
-6 -5
32
-4
-3
33
SURFACE [4105] {d}
TIME= 900000
-2
-1
0
34
1
2
35
3 -6 -5
32
-4
-3
33
SURFACE [4106] {d}
TIME= 1100000
-2
-1
0
34
1
2
35
3 -6 -5
32
-4
-3
33
SURFACE [4107] {d}
TIME= 1200000
-2
-1
0
34
1
2
35
3 -6 -5
32
-4
-3
33
SURFACE [4108] {d}
TIME= 1300000
-2
-1
0
34
1
2
35
3 -6 -5
32
-4
-3
33
TIME= 1400000
-2
-1
0
34
1
2
35
3 -6 -5
32
-4
-3
33
-2
-1
0
34
1
2
35
3
43
43
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43
43
43
43
42
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42
42
42
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42
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41
41
41
41
41
41
40
40
32
33
34
35
40
32
THICKNESS [4102] {d}
33
34
35
THICKNESS [4103] {d}
TIME= 800000
40
32
33
34
35
THICKNESS [4105] {d}
TIME= 900000
40
32
33
34
35
THICKNESS [4106] {d}
TIME= 1100000
40
32
33
34
35
THICKNESS [4107] {d}
TIME= 1200000
40
32
33
34
35
THICKNESS [4108] {d}
TIME= 1300000
TIME= 1400000
0.0 0.5 1.0 1.5 2.0 2.5 3.0 3.5 4.0 4.5 5.0 5.5 0.0 0.5 1.0 1.5 2.0 2.5 3.0 3.5 4.0 4.5 5.0 5.5 0.0 0.5 1.0 1.5 2.0 2.5 3.0 3.5 4.0 4.5 5.0 5.5 0.0 0.5 1.0 1.5 2.0 2.5 3.0 3.5 4.0 4.5 5.0 5.5 0.0 0.5 1.0 1.5 2.0 2.5 3.0 3.5 4.0 4.5 5.0 5.5 0.0 0.5 1.0 1.5 2.0 2.5 3.0 3.5 4.0 4.5 5.0 5.5
32
33
34
35
32
33
34
35
32
33
34
35
32
33
34
35
32
33
34
35
32
33
34
35
43
43
43
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43
43
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42
42
42
42
41
41
41
41
41
41
41
40
40
40
32
33
34
32
35
VELOCITY [4102] {d}
33
34
35
VELOCITY [4103] {d}
TIME= 800000
33
34
32
35
VELOCITY [4105] {d}
TIME= 900000
-10 -9 -8 -7 -6 -5 -4 -3 -2 -1 0 1 2
32
33
35
34
40
40
32
34
35
40
32
VELOCITY [4106] {d}
TIME= 1100000
3 -10 -9 -8 -7 -6 -5 -4 -3 -2 -1 0 1 2
33
35
32
34
33
34
35
40
32
VELOCITY [4107] {d}
TIME= 1200000
3 -10 -9 -8 -7 -6 -5 -4 -3 -2 -1 0 1 2
33
35
32
34
33
34
35
VELOCITY [4108] {d}
TIME= 1300000
3 -10 -9 -8 -7 -6 -5 -4 -3 -2 -1 0 1 2
32
33
34
35
33
TIME= 1400000
3 -10 -9 -8 -7 -6 -5 -4 -3 -2 -1 0 1 2
35
32
33
34
3 -10 -9 -8 -7 -6 -5 -4 -3 -2 -1 0 1 2
35
32
33
34
3
43
43
43
43
43
43
43
42
42
42
42
42
42
42
41
41
41
41
41
41
41
40
40
32
33
34
35
40
32
33
34
35
40
32
33
34
35
40
32
33
34
35
40
32
33
34
35
40
32
33
34
35
Fig. 12. Surface (m), thickness (m), and velocity (log 10(mm/a)) at 800, 900, 1100, 1200, 1300, and 1400 Ka in the DPM lineated valley fill system.
for example, if the source region in the Tharsis region becomes depleted, or if the obliquity changed (which is to be expected on
approximately this time scale; Laskar et al., 2004). We remove
the positive accumulation component by separating the positive
(accumulation) and negative (ablation) portions of the SMB provided by the GCM results.
A typical modeled point from the GCM results is shown in
Fig. 11. The amount of ice equivalent on the surface is shown
as a function of time throughout the martian year. Ice amount
begins at zero and climbs to close to 3 kg/m2 (a period of positive accumulation). Ice then begins to thin (negative ablation).
We allowed the ice amount to go negative in the LMD/GCM to
have access to the amount of ice that would be ablated, if there
were any there to ablate. After this period of sublimation, a second period of positive accumulation occurs, ending the year with
a net accumulation of 10 kg/m2. It is this number that is reported in Fig. 8a. In this analysis, on the other hand, we keep
track of the separate amounts of positive accumulation (ice
growing) and negative ablation (ice thinning). Our difference between these two components of the SMB corresponds exactly to
the values reported by Madeleine et al. (2009) in Fig. 8a, but we
have a measure of potential negative mass balance (a point
where the final endpoint of ice amount was negative) and we
have the possibility to modify either the positive part (turn off
the snowfall) or the negative part (increase or decrease sublimation) while keeping the rest of the climate true to the GCM results (the spatial distribution of mass balance and surface
temperatures).
We now explore the retreat of the regional ice sheet. In this
model we do not include any effects of debris cover that might
be produced as the valley and dichotomy boundary scarps become exposed and could potentially provide a rock and debris
cover to armor the ice and protect it from sublimation. Instead,
the retreating ice is assumed to be pure ice. The configuration
of the DPM valley at 600 Ka, the point at which retreat is allowed to begin by removing the positive portion of the mass
balance, is shown in the right-most column of Fig. 10. The evolution of the ice sheet as this retreat progresses is shown in
Fig. 12.
(1) By 800 Ka (Fig. 12, first column), 200 Ka after retreat begins,
the surface is dropping and the underlying topography is
becoming more evident as the first aretes emerge as nunataks. Velocities are considerably reduced to 1 mm/a in the
valleys and 0.001 mm/a on the plateaus, although still
clearly channelized in the valley proper.
J.L. Fastook et al. / Icarus 216 (2011) 23–39
(2) As ice sheet collapse continues, the plateaus begin to emerge
by 900 Ka (Fig. 12, second column), and are fully cleared by
1100 Ka (Fig. 12, third column).
(3) By 1200 Ka (Fig. 12, fourth column), ice in the valleys is less
than 500 m thick, while on the adjacent plains ice thickness
is still over 1 km.
(4) By 1300 Ka (Fig. 12, fifth column), the remaining patches of
ice separate, and the valleys are ice-free (again, assuming no
debris cover).
(5) Of particular interest is the configuration at 1400 Ka (Fig. 12,
right-most column), where only limited patches remain surrounding nunataks on the plains outside the mouth of the
valley, a configuration similar to the patterns for lobate debris aprons reported by Ostrach and Head (2008), Ostrach
et al. (2008), Holt et al. (2008a,b), Plaut et al. (2008), Safaeinili et al. (2009) and Kress and Head (2009). These relationships predict that the lobate debris aprons (LDA) might
slightly postdate the lineated valley fill (LVF) in their formation. These remnant patches occur along the base of steep
topographic scarps because the ice is thickest where it flows
from the high to lower terrain across a step in bed elevation.
While not modeled in this experiment, the potential for debris armoring beneath scarps would increase the likelihood of
LDA preservation.
In summary, coupling the regional accumulation rates of the
LMD/GCM (Madeleine et al., 2009) with a glacial flow model shows
that the resulting regional plateau and plains glaciation that would
have characterized the broad area during the period of glaciation is
quite capable of producing results during its collapse that are consistent with the alpine valley glaciation interpretation of the LVF.
Mapping of the patterns of retreat during the waning stages of
the glaciation shows that the valley walls are eventually exposed
and could readily serve as sources of debris to produce the lineated
valley fill and related flow patterns seen in numerous places along
the dichotomy boundary (Fig. 1a–e). Furthermore, the final configuration of the ice sheet upon its retreat is that of glacial ice surrounding massifs in the plains just north of the dichotomy
boundary, a pattern very similar to that of lobate debris aprons
that are currently observed to surround the massifs. We are currently using the glacial flow models illustrated in Fig. 12 to model
the exposure of the rock cliffs and massifs and assess the effects
that the addition of cliff-derived rock debris will have on the sublimation rates and flow rates of the lineated valley fill and lobate
debris aprons (Fastook et al., 2010; e.g., Kowalewski et al., 2011).
37
In the second glacial flow model simulation, the regional accumulation patterns predicted by the LMD/GCM were used (Madeleine et al., 2009). This glacial flow model simulation produced a
much wider region of thick ice accumulation, and significant glaciation on the plateaus and in the regional plains surrounding the
dichotomy boundary (Figs. 1a and 8a). Deglaciation produced
decreasing ice thicknesses, with flow centered on the fretted valleys. As plateaus lost ice, and scarps and cliffs of the valley and
dichotomy boundary walls were exposed, there was considerable
potential for the production of a rock debris-cover that could preserve the underlying ice and the surface flow patterns seen today.
In this model, the lineated valley fill and lobate debris aprons were
the product of final retreat and downwasting of a much more regional glacial landsystem (Marchant and Head, 2008), rather than
the maximum extent of an alpine valley glacial landsystem.
From these glacial flow models, information can be derived
about the possible timing of the growth of these systems. In the alcove-only case, several millions of years are required for the flow
system to achieve a pattern consistent with the geological record.
Because the alcove-only case uses specified areal distribution of ice
deposition we cannot constrain its occurrence to any specific time
interval.
Coupling the ice sheet model with a GCM yields a pattern of ice
sheet development, which during its collapse also yields a configuration consistent with the geological observations. Because of
the high accumulation rates predicted by the LMD/GCM, the ice
sheet grows much more rapidly than the alcove-only case. As the
ice sheet collapses, presumably due to the exhaustion of the moisture source in the Tharsis region or spin-axis/orbital changes, a
configuration emerges of a limited valley glacier system compatible with the geological observations. Importantly, as the surface
lowers to the point where armoring debris would be available from
the cliff scarps, the configuration resembles the alcove-only case.
At this point, if debris is supplied from the adjacent newly exposed
rock cliffs, we would expect the sublimation rates over the debriscovered valley-floor glaciers to be reduced by orders of magnitude
(Kowalewski et al., 2006) potentially preserving ice even now in
the valley trunks. This will also be the case with isolated massifs
at the mouths of the valleys, leading to the possibility that the lobate debris aprons are ice-cored remnants of a much more extensive ice sheet that covered the area during climatically appropriate
periods. We are currently exploring these options by modeling the
accumulation of debris once cliffs are exposed, and the evolution
and flow of such debris-covered glaciers (Fastook et al., 2010).
Acknowledgments
4. Conclusions
A fretted valley system located at the northern mid-latitude
dichotomy boundary and containing lineated valley fill with extensive flow-like features interpreted to be glacial in origin (Head
et al., 2006a) has been modeled using glacial flow models linked
to atmospheric general circulation models. The glacial flow model
results are consistent with the geological features observed
and with the interpretation of these as alpine valley glacial
landsystems.
In the first glacial flow model simulation, sources were modeled
in the alcoves only and were found to be consistent with the alpine
valley glaciation interpretation for various environments of flow in
the system. These results indicated that the currently observed LVF
deposits could be the result of initial ice accumulation in the alcoves, accompanied by debris cover that led to advancing alpine
glacial landsystems to the extent observed today, with preservation of their flow texture and the underlying ice during downwasting in the waning stages of glaciation.
We gratefully acknowledge the scientists and engineers contributing to the success of the NASA Mars Observer, Mars Odyssey,
and Mars Reconnaissance Orbiter, and the ESA Mars Express missions. JWH gratefully acknowledges support from NASA through
the Mars Data Analysis Program and the Mars Express HRSC Team.
J.L.F. gratefully acknowledges support from the NSF for development of the terrestrial version of UMISM through many grants.
D.R.M. and J.W.H. gratefully acknowledge support from the NSF
through Grant ANT-0944702. Thanks are extended to Anne Côté
for assistance in manuscript preparation.
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