Icarus 237 (2014) 315–339
Contents lists available at ScienceDirect
Icarus
journal homepage: www.elsevier.com/locate/icarus
Volcano–ice interactions in the Arsia Mons tropical mountain glacier
deposits
Kathleen E. Scanlon a,⇑, James W. Head a, Lionel Wilson b, David R. Marchant c
a
Department of Geological Sciences, Brown University, Providence, RI 02912, USA
Lancaster Environment Centre, Lancaster University, Lancaster LA1 4YQ, UK
c
Department of Earth and Environment, Boston University, Boston, MA 02215, USA
b
a r t i c l e
i n f o
Article history:
Received 7 February 2014
Revised 11 April 2014
Accepted 16 April 2014
Available online 26 April 2014
Keywords:
Mars
Mars, surface
Volcanism
a b s t r a c t
Fan-shaped deposits (FSD) superposed on the sides of the Tharsis Montes volcanic edifices are widely
interpreted to have been formed by cold-based glaciation during the Late Amazonian, a period when
the Tharsis Montes were volcanically active. We survey the 166,000 km2 Arsia Mons FSD using new,
high-resolution image and topography data and describe numerous landforms indicative of volcano–
ice interactions. These include (1) steep-sided mounds, morphologically similar to terrestrial tindar that
form by subglacial eruptions under low confining pressure; (2) steep-sided, leveed flow-like landforms
with depressed centers, interpreted to be subglacial lava flows with chilled margins; (3) digitate flows
that we interpret as having resulted from lava flow interaction with glacial ice at the upslope margin
of the glacier; (4) a plateau with the steep sides and smooth capping flow of a basaltic tuya, a class of
feature formed when subglacial eruptions persist long enough to melt through the overlying ice; and
(5) low, areally extensive mounds that we interpret as effusions of pillow lava, formed by subglacial eruptions under high confining pressure. Together, these eruptions involved hundreds of cubic kilometers of
subglacially erupted lava; thermodynamic relationships indicate that this amount of lava would have
produced a similar volume of subglacial liquid meltwater, some of which carved fluvial features in the
FSD. Landforms in the FSD also suggest that glaciovolcanic heat transfer induced local wet-based flow
in some parts of the glacier. Glaciovolcanic environments are important microbial habitats on Earth,
and the evidence for widespread liquid water in the Amazonian-aged Arsia Mons FSD makes it one of
the most recent potentially habitable environments on Mars. Such environments could have provided
refugia for any life that developed on Mars and survived on its surface until the Amazonian.
Ó 2014 Elsevier Inc. All rights reserved.
1. Introduction
Volcanism is known to be a fundamental process in shaping the
surface of Mars throughout its history, including the early volcanic
construction of Tharsis, the widespread late Noachian and Early
Hesperian ridged plains, and Amazonian volcanism of the Tharsis
and Elysium region (e.g., Tanaka et al., 1991; Head, 2007;
Werner, 2009; Carr and Head, 2010a). Historically, water ice on
Mars has been thought to reside primarily in the regolith (e.g.,
Mellon and Jakosky, 1993; Mellon et al., 2004), upper crust
cryosphere (e.g., Clifford, 1993), and thick polar ice deposits (e.g.,
Thomas et al., 1992; for extensive reviews see Carr, 1996; Byrne,
2009; Carr and Head, 2010b). In the mid-1990s, increasing
numbers of workers began to appreciate the fact that past spin
⇑ Corresponding author.
E-mail address: kathleen_scanlon@brown.edu (K.E. Scanlon).
http://dx.doi.org/10.1016/j.icarus.2014.04.024
0019-1035/Ó 2014 Elsevier Inc. All rights reserved.
axis–orbital parameter variations could cause polar ice to be mobilized, transported and deposited equatorward (as pointed out by
Jakosky and Carr, 1985), raising the possibility that volcanic eruptions might interact with deposits of ice, not just in the subsurface
(Chapman and Tanaka, 2002), but also in glacial (Head and Wilson,
2002) and frozen lake environments (Chapman, 1994; Chapman
and Tanaka, 2001). The development of analytical solutions for
the possible histories of spin axis/orbital parameters on Mars
(Laskar et al., 2004) paved the way for an increased understanding
of locations where non-polar ice might reside under different past
conditions (e.g., Head et al., 2003; Head and Marchant, 2003; Shean
et al., 2005; Baker et al., 2010) and the climatic and atmospheric
settings in which they might have been deposited (e.g., Forget
et al., 2006; Madeleine et al., 2009, 2013). Concurrently, interest
in astrobiology has led to a search for specific environments that
might have been conducive to initiation and sustenance of life
(The MEPAG Special Regions–Science Analysis Group, 2006),
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K.E. Scanlon et al. / Icarus 237 (2014) 315–339
including microbial habitats related to volcano–ice interactions
(Cousins and Crawford, 2011).
New high-resolution image, spectral, geochemical and altimetry
data collected by recent spacecraft missions have provided the
ability to analyze the interaction of volcanic intrusions and eruptions with widespread paleo-ice deposits in unprecedented detail.
In this contribution, we analyze the Arsia Mons fan-shaped deposit
(FSD), a 166,000 km2 deposit on the northwest flank of the Arsia
Mons shield volcano interpreted to have formed in the Late Amazonian as a tropical mountain glacier (Head and Marchant, 2003;
Forget et al., 2006; Shean et al., 2005, 2007; Fastook et al., 2008).
We review the properties of the FSD and the nature and history
of Tharsis volcanism, outline the basic principles of volcano–ice
interactions, and examine specific examples of volcano–ice interactions in the Arsia Mons deposit. We find abundant examples of
glaciovolcanic landforms in the Arsia Mons deposits.
2. Volcanism, glaciation and volcano–ice interactions in the
Tharsis region
2.1. Volcanism in the Tharsis region
Volcanic activity has been key in forming the Tharsis rise (e.g.
Phillips et al., 2001) as well as the individual major volcanic edifices that characterize its surface: Olympus Mons, Alba Patera,
and the Tharsis Montes, Ascraeus, Pavonis and Arsia Mons. For
Arsia Mons in particular, associated lava flows date from Middle
Hesperian to Late Amazonian in age, with major episodes of volcanism ending at 100–200 Ma, 500 Ma, 800 Ma and 2 Ga and
major shield construction ending ca. 3.54 Ga (Werner, 2009). Based
on crater counts and stratigraphic relationships, Crumpler and
Aubele (1978) suggested that volcanism at Arsia Mons followed a
sequence in which (1) thick lava plains accumulated throughout
the Tharsis region, followed by (2) construction of the main shield
by short multiple flow units similar to those that build terrestrial
shields; (3) the onset of intense eruptive activity forming ‘‘parasitic
shields’’ along a southwest to northeast trending linear rift or fissure; (4) formation of concentric graben, concentric fractures,
and the caldera, possibly contemporaneously with (3); and (5)
continuing eruptions on the northeast and southwest flanks.
Bleacher et al. (2007) noted evolution towards shorter flow lengths
and from tube-fed flows to channel-fed flows over the history of
each of the Tharsis Montes; terrestrial basaltic tube-fed flows are
associated with lower-viscosity lavas, lower effusion rates and
longer-term, more stable eruptions than their channel-fed
counterparts.
Increasing silica content decreases the capacity of lava to melt
ice and increases the aspect ratio of glaciovolcanic landforms
(Hickson, 2000; Smellie, 2007; McGarvie, 2009). The composition
of Tharsis lavas has been difficult to discern, however, because
the thick mantle of dust covering the region (Ruff and
Christensen, 2002) prevents definitive spectroscopic measurements of mineralogy. Linear deconvolution of TES spectra from relatively dust-free flows near Arsia, predating rift apron formation,
indicated that the flows have a basaltic composition similar to
martian shergottite meteorites (Lang et al., 2009). A basaltic composition can also be inferred for some Tharsis flows based on the
distance they traveled from their source vents (Scott and
Zimbelman, 1995; Hiesinger et al., 2007).
2.2. Glaciation in the Tharsis region
The FSDs on the west and northwest flanks of the Tharsis Montes consist primarily of three units (Zimbelman and Edgett, 1992;
Scott and Zimbelman, 1995): a ridged facies, a knobby facies, and
a smooth facies that is uniformly stratigraphically higher than
the other two. Previously proposed hypotheses have postulated a
wide variety of candidate origins for the FSDs, including masswasting processes, pyroclastic flows or other volcanic processes,
tectonic deformation, or glaciation (see review in Shean et al.,
2005). Williams (1978) was the first to propose that the deposits
might be glacial. Lucchitta (1981) noted the similarities between
ridges in the Olympus and Arsia FSDs and recessional glacial ridges
in Alaska and Iceland, and suggested that the even superimposition
of some martian ridges upon high topographic obstacles implied
the removal of a large quantity of material, most easily explained
by the sublimation of ice.
Head and Marchant (2003) further refined this explanation of
the ridged facies by drawing an analogy to the ridge-forming drop
moraines left by cold-based glaciers in the Mars-like Antarctic Dry
Valleys. Based on its superposition relationships, its association
with the ridged facies, and its lack of fluvial features, they argued
that the knobby facies represents sublimation till; based on its
morphology and spatial associations, they argued that the smooth
facies consists of remnant debris-covered glaciers. Shean et al.
(2005), using data from the Thermal Emission Imaging System
(THEMIS), the Mars Orbital Laser Altimeter (MOLA), and the Mars
Orbiter Camera (MOC) in addition to Viking imagery, mapped
glacial and glaciovolcanic features in the Pavonis Mons FSD,
including tindar, outflow channels, and lava flows with steep,
ice-chilled margins. Shean et al. (2005) presented counterarguments to each of the problems with the glacial hypothesis for
Tharsis FSD formation raised by Edgett (1989). Kadish et al.
(2008) mapped a similar array of features in the Ascraeus Mons
FSD, as well as a prominent plateau resembling terrestrial subglacially erupted tuyas. Shean et al. (2007) identified a younger
(65 Ma), smaller set of cold-based glacial deposits, including
remnant debris-covered glaciers, superimposed on the Arsia Mons
fan-shaped deposit.
Climate reconstructions and results from ice-flow modeling
have lent additional support to the geologic arguments that the
Tharsis FSDs are glacial in origin. Simplified ice stability models
have predicted for decades (Jakosky and Carr, 1985; Jakosky
et al., 1995) that water vapor would be transported from the poles
of Mars to lower latitudes during periods of high spin-axis obliquity, though these models were unable to predict the resulting ice
distribution in detail. More recently, three-dimensional general
circulation models (GCMs) that include water cycle physics have
consistently simulated the process (Haberle et al., 2000;
Mischna et al., 2003; Levrard et al., 2004). The GCM experiments
of Forget et al. (2006) were the first with sufficient resolution to
allow direct comparison with geological features. In their simulations at 40° and higher obliquity, sublimation of polar ice
increased the atmospheric water content relative to the current
martian climate. As this moist air ascended the Tharsis Montes
and Olympus Mons in the prevailing westerly to northwesterly
winds, precipitation occurred on the windward flanks of the volcanoes. These results further supported the interpretation of the
Tharsis FSDs as tropical mountain glaciers (Head and Marchant,
2003).
In the Forget et al. (2006) GCM simulations, ice precipitated at
the upslope region of the deposit, where the accumulation zone
of the hypothesized glacier would be expected. Fastook et al.
(2008) used climate fields from these simulations as input in
model runs using the University of Maine Ice Sheet Model
(UMISM). By parameterizing mass balance in terms of local temperature and restricting the orientation of the long axis of glacial
flow, the Arsia Mons ice sheet reproduced in model runs matched
the footprint of the mapped Arsia Mons FSD with an accuracy comparable to that of terrestrial models in simulating modern ice
sheets (Fastook et al., 2008). In summary, the current consensus
K.E. Scanlon et al. / Icarus 237 (2014) 315–339
favors the glacial hypothesis for the origin of the Tharsis FSDs
(Williams, 1978; Lucchitta, 1981; Head and Marchant, 2003). We
thus proceed to investigate the interaction of volcanism and glaciation in the Tharsis region, with particular emphasis on the Arsia
Mons fan-shaped deposit.
2.3. Principles of volcano–ice interactions
The interaction of ascending magma and glacial ice involves
significant heat transfer and melting, and thus the process can
have a profound influence on the nature and fate of the magma,
the properties of the glacial ice, the resulting volumes and fate
of meltwater produced, and the resulting landforms, both glacial
and volcanic. Several terrestrial environments (e.g., Iceland, the
Antarctic Peninsula, and western Canada) have proven to be natural laboratories for the study of these interactions and resulting
landforms and for the development of the basic physical
principles of heat transfer and melting (e.g., Smellie and Skilling,
1994; Gudmundsson, 1997; Hickson, 2000). We use the series of
principles and physical processes related to magma–ice interactions in terrestrial glaciers developed by Höskuldsson and
Sparks (1997), Wilson and Head (2002, 2007a, 2009),
Gudmundsson (2003), Head and Wilson (2007), Tuffen (2007),
and Wilson et al. (2013), and applied to Mars by Head and
Wilson (2002, 2007), to assess the relationships in the Arsia Mons
fan-shaped deposit.
2.4. Volcano–ice interactions and microbial habitats
Interaction between glaciers and lava flows on Earth creates
environments that can function as microbial habitats, including
subglacial lakes, subglacial basaltic edifices, permafrost hydrothermal systems, and glacial springs (Cousins and Crawford,
2011). Volcano–ice interactions can also induce locally wet-based
conditions in cold-based glaciers (e.g., Fahnestock et al., 2001; Fox
Maule et al., 2005; Head and Wilson, 2007; Hambrey et al., 2008),
which introduce new habitable environments including wet subglacial sediments. Wet-based glacial ice also contains more
water-filled fractures above the bed (Smellie, 2006), which can
serve both as habitats and as pathways for any potential life to
better distribute itself throughout a glacier (Fountain et al.,
2005). In this work, we document evidence for many episodes
of volcano–ice interaction in the Arsia Mons FSD; Scanlon et al.
(in preparation, 2014) document the effects of this heat transfer
on the style of glaciation as recorded in the deposit, and the
potential for habitable environments in the Late Amazonian Arsia
Mons FSD glacier.
2.5. Data and methods
A photographic basemap covering Arsia Mons and the Arsia fanshaped deposit was constructed from Mars Reconnaissance Orbiter
(MRO) Context Camera (CTX) images (Malin et al., 2007) with
5 m per pixel resolution using ArcMap 10.0. This basemap was
augmented with images from the High Resolution Stereo Camera
(HRSC) with 10–30 m per pixel resolution (Neukum and
Jaumann, 2004). Topographic data from MOLA (Zuber et al.,
1992; Smith et al., 1999) at 128 pixel per degree (463 m per
pixel) resolution and, where available, HRSC-derived Digital Elevation Maps (DEMs) with 100 m per pixel resolution (Dumke et al.,
2008), were superimposed on the CTX basemap. These data were
used with the Spatial Analyst toolkit in ArcMap 10.0 to create slope
maps, contour maps, elevation profiles, and other derived products. The Arsia FSD and its surroundings (Fig. 1) were examined
and mapped in detail.
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3. Candidate glacial and glaciovolcanic features of the Arsia
Mons fan-shaped deposit
Zimbelman and Edgett (1992), Scott and Zimbelman (1995),
and Head and Marchant (2003) describe the geological setting
and the main units of the Arsia Mons FSD. Shean et al. (2007)
present a study of recent glaciation on Arsia Mons. Here we catalog
glacial and glaciovolcanic features of the Arsia Mons FSD (Fig. 1),
including those previously noted by Wilson and Head (2007b,c)
but not described in detail (Fig. 2), and present a comprehensive
examination of candidate glaciovolcanic features.
3.1. Northwest plateau
Toward the northwest edge of the deposit (Fig. 1b, box 3.1) lies
a plateau 17 km long by 15 km wide and 140 m high (Fig. 2a
and b). The plateau is steep-sided compared with nearby subaerial
lava flows, having typical slopes between 6° and 12° (Fig. 2c). An
elongate, southwest–northeast oriented mound, 10 km long and
3 km wide, stands at the center of the plateau. The mound slopes
steeply (angles up to 15°) to the northwest and slightly more
gently to the southeast, and stands 150 m (up to 200 m) above
the rest of the plateau (Fig. 3). At its northern edge, the plateau
conforms closely to and becomes indistinguishable from the rim
of a heavily degraded crater 3.5 km in diameter. The circularity
and rim structure of the crater distinguish it from volcanic pit craters found elsewhere in the deposit, and we interpret it as an
impact crater predating the formation of the plateau. The plateau
lies along the trend of a line of subaerial spatter cones beginning
45 km to the north (Fig. 4) (Head and Wilson, 2007), suggesting
a possible genetic relationship along the same dike-emplacement
strike.
Superposed on the plateau are two sinuous ridges (Fig. 5a and
b), originating near the base of the mound. The southern ridge is
4–6 m tall and extends at least 25 km downslope; the northern
ridge is up to 20 m tall but becomes difficult to trace 5 km
beyond the edge of the plateau. Both ridges have sharply defined
flat crests atop the plateau (Fig. 5c) and take on a more diffuse
appearance downslope. Downslope (i.e., to the northwest) of the
plateau, unusually tall and thick moraines are bowed outwards
relative to the surrounding drop moraines (Fig. 4a and b). Knobs
across a region downslope of the plateau appear aligned or elongated in the downslope direction (Fig. 4).
More than a dozen channels emerge from beneath the prominent moraine downslope of the plateau (Fig. 4a and b); 6 km further downslope the channels coalesce into a braided channel
which is traceable for 35 km (Fig. 6a and b). The channels emerge
from beneath a layer of fine sediment at the base of the moraine.
Where HRSC DEM coverage is available, we determined that the
main channel is 35 m deep and has a V-shaped profile (Fig. 7a);
the tributaries are shallower (Fig. 7b).
Calling on the series of principles and physical processes developed by Wilson and Head (2002, 2007a, 2009), Head and Wilson
(2007), Tuffen (2007), and Wilson et al. (2013), and applied to Mars
by Head and Wilson (2002, 2007; Fig. 8), we interpret this collection of features to be related to the interaction of one or more dike
emplacement events with the Arsia tropical mountain glacier, and
the aftermath of the contact of hot intruding magma with the overlying glacial ice.
When the confining pressure of ice above a subglacial eruption is
sufficient to suppress explosive degassing, the lava emerges as pillows (e.g., Wilson and Head, 2002; Tuffen, 2007). The pillows form
low mounds, ridges, or sheets that slope gently when emplaced into
water, but can become much steeper if they chill against ice walls
(Smellie, 2009). Meltwater produced by the eruption is confined
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K.E. Scanlon et al. / Icarus 237 (2014) 315–339
Fig. 1. (a) Geomorphological map of the Arsia Mons fan-shaped deposit (FSD), after Zimbelman and Edgett (1992) and Scott and Zimbelman (1995). Red lines denote large
volcanic graben, white lines denote contacts between units, and black lines denote the outlines of glaciovolcanic landforms and distorted moraines. Dashed lines denote
inferred contacts. Blue and white dots in the smooth facies denote pits and knobs, respectively. THEMIS 100 m/pixel daytime image mosaic. (b) Numbered white boxes
indicate the regions discussed in Sections 3.1–3.7 of this paper. All map figures in this paper are oriented north upwards. (For interpretation of the references to color in this
figure legend, the reader is referred to the web version of this article.)
K.E. Scanlon et al. / Icarus 237 (2014) 315–339
319
Fig. 2. The Northwest plateau; (a) HRSC DEM superimposed on CTX image mosaic. (b) Sketch map. The ridged facies is shaded in red and the knobby facies is shaded in blue.
Distorted drop moraines are denoted by dashed lines, fluvial channels by light blue lines, grooves or channels by light gray lines, and streamlined knobs by green lines. The
transects of the profiles in Figs. 3 and 5c are shown as thick white lines. (c) Slope map derived from an HRSC DEM. Contacts between the ridged and knobby facies are shown
by white lines; other geological features are outlined in black. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of
this article.)
to a cavity or cavities in the glacier, above the flow (Fig. 8a). If the
pressure in the melt cavity is reduced sufficiently by the lava flow
approaching the edge of the glacier, meltwater draining from under
the glacier, or the growing edifice thinning the ice above it by melting, the eruption can become explosive (Wilson and Head, 2002;
Tuffen, 2007; Wilson et al., 2013). Explosive eruptions melt ice
much more efficiently than pillow lava effusion because the lava
fragmentation creates a much larger surface area per unit volume
of lava (Head and Wilson, 2002; Tuffen, 2007), and the melt cavity
becomes much larger immediately above the vent, growing outward as the eruption continues (Fig. 8b).
On the basis of morphology, morphometry, and landform associations (Figs. 5–7), and physical volcanology elucidated by terrestrial subglacial eruptions (Fig. 8), we interpret the history of the
Northwest plateau complex as follows (Fig. 9):
1. Lava erupted under thick ice cover and spread outward and
downslope as pillows, embaying a nearby, pre-glacial, impact
crater (Fig. 9a). This flat, steep deposit formed the majority of
the Northwest plateau. While it is larger than any known terrestrial examples, the main body of the Northwest plateau lies on
the area-to-volume trend line of terrestrial pillow mounds and
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K.E. Scanlon et al. / Icarus 237 (2014) 315–339
Fig. 3. (a) Linear detrended profile across the line A–A0 in Fig. 2b. (b) Linear detrended profile across the line B–B0 in Fig. 2b. (c) Linear detrended profile across the line C–C0 in
Fig. 2b. (d) Linear detrended profile across the line D–D0 in Fig. 2b. Data from HRSC-derived DEM.
sheet-like sequences (Smellie, 2009, Fig. 6). The water melted
by the pillow lava effusion remained confined in a lens above
the flow (Fig. 9a).
2. Elevated geothermal heat in the vicinity of the eruption, or lifting of the glacier by the growing melt body (Smellie, 2000), may
have caused the glacier to decouple from the ground (Fig. 9b).
This layer of water would have been in contact with the melt
cavity, at glaciostatic pressure. Due to the decoupling, the glacier transitioned from regionally cold-based to locally wetbased and was then able to slide along its bed, sculpting knobs
downslope of the plateau into elongated forms (Fig. 9a and b).
This caused a lobe to extend farther downslope than the surrounding (cold-based) glacier, as evidenced by the downslope
bulge of the glacial drop moraines in the vicinity of the plateau:
while the moraines are a recessional feature that would not
have been deposited until after the eruption ceased, the local
deformation of the ice in this region would have changed the
local outline of the glacier, and hence the shape of the drop
moraines left behind as it receded.
3. Once the layer of liquid water reached the edge of the glacier,
the pressure difference between the subglacial environment
and the atmospheric pressure outside the glacier caused the
water to flow out catastrophically in a jökulhlaup (Fig. 9c); a
transition from subglacial to subaerial flow is indicated by (a)
the braided channel morphology outside the moraine we interpret as marking the syneruptive glacial margin and (b) the
many tributary heads emerging from this moraine. This outflow
carved the channels downslope (Figs. 5–7).
4. The drainage of the melt cavity suddenly reduced the water
pressure (and hence the pressure above the vent) from glaciostatic to atmospheric; the eruption was no longer constrained
to a pillow effusion. Since volatiles were then free to exsolve
and water could still enter the vent, the eruption entered an
explosive phase (Fig. 9d), melting a taller cavity into the ice
above the vent (Wilson and Head, 2002; Wilson et al., 2013).
The tephra from this eruption, confined in the new cavity,
formed the mound over the center of the plateau.
5. Melt continued to escape through the jökulhlaup channels until
the collapse or deformation of the ice under its own weight
closed the thin cavity over the pillow sheet. Due to the creation
of the new melt cavity, and the downward motion of the ice
surface above it caused by density differences between ice
and water (Smellie, 2000), the ice above the vent thinned considerably and the eruption could proceed as explosive until
the magma supply was exhausted.
6. As the edifice cooled, some of the remaining meltwater eroded
channels into the ice at the base of the glacier (Fig. 9e). More and
coarser hyaloclastite tephra grains fell out of suspension at the
heads of the channels. The sediment in the melt channels formed
the present-day eskers, which are tall and sharp-crested close to
the vent, but low and broad-crested downslope (Figs. 5–7).
Sharp-crested and broad-crested eskers are left where subglacial
channels rapidly melted away the surrounding ice and where
channel water froze onto the channel walls, respectively (e.g.
Shreve, 1985). This transition from sharp- to broad-crested
eskers away from the plateau therefore supports the hypothesis
that meltwater in the channels was generated by heat from the
volcanic edifice.
3.2. Elongate plateau
Centered 45 km north of the Northwest plateau (Fig. 1b, box
3.2) lies an elongate plateau 39 km long, with a maximum width
of 18 km (Fig. 10a and b). The plateau trends approximately
K.E. Scanlon et al. / Icarus 237 (2014) 315–339
321
Fig. 4. (a) THEMIS daytime IR context image of the region surrounding the northwest and elongate plateaus. (b) Sketch map of the same region. Channels are outlined in blue,
chains of pit craters or spatter cones are outlined in violet, and a lava flow bifurcated by the outermost moraine of the FSD is outlined in red. The ridged facies is shaded in
light gray and the knobby facies is shaded in dark gray. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this
article.)
north–south and has a height of up to 140 m. Irregular lobes
extend downslope from its center and a gently sloping lobe
extends downslope from its northern terminus (1–6°, compared
with 12° or steeper slopes on other edges of the plateau;
Figs. 10c and 11b). At its center lies an elongate, west–southwest
to east–northeast trending mound which stands up to 140 m
higher than the rest of the plateau (Figs. 10a, b and 11a, c).
Several overlapping, highly sinuous ridges extend 5–20 km
north and west from the central mound of the plateau. As with
the Northwest plateau, many drop moraines near the plateau
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K.E. Scanlon et al. / Icarus 237 (2014) 315–339
Fig. 6. Fluvial channels near the Northwest plateau. (a) CTX image of channels
emerging from the Arsia Mons fan-shaped deposit. The channels appear to incise a
layer of subglacial sediment (white arrows) as they emerge from the glacier, and are
superposed by concentric, regular moraines (black arrows). The topographic profile
across the line A–A0 is shown in Fig. 7a; the topographic profile across the line B–B0
is shown in Fig. 7b. (b) Sketch map of the channels shown in (a).
Fig. 5. Ridges, interpreted to be eskers, superimposed on the Northwest plateau. (a)
Close view of eskers superimposed on the Northwest plateau and streamlined
boulders to the north. CTX image mosaic. (b) Sketch map of the region in (a). The
ridged facies is shaded in red and the knobby facies is shaded in blue. (c) Profile
across an esker, along the line E–E0 in Fig. 2b. Data from HRSC-derived DEM. (For
interpretation of the references to color in this figure legend, the reader is referred
to the web version of this article.)
bow outwards relative to superposed and nearby drop moraines
(Fig. 10a and b). The moraine that we interpret as representing
the glacier boundary at the time of plateau formation, due to its
enlarged and knobby appearance and stratigraphic relationships
with the other local landforms, is also bent outwards.
A gently sinuous channel, sharply incised at the edge of the plateau and surrounded by smooth debris farther downslope, begins
near the northern terminus of the plateau (Fig. 12). It is superposed
by the large moraine that we interpret as having been deposited by
the syneruptive glacial margin, and crosscuts several drop moraines farther downslope (Fig. 12). The contact between the flows
emerging from the line of spatter cones west of the plateau and
the debris associated with the channel is not visible at CTX resolution, but the edges of the flows are sharply defined whereas the
boundaries of the debris are diffuse, and the surface texture of
the flows is knobby compared with that of the channeled debris.
The long axis of the plateau is collinear with a curving graben to
the south and a series of pits to the north (Fig. 4a and b). The pits
cross-cut a ridged lava flow; this flow, in turn, is bifurcated by the
outermost moraine of the FSD (Fig. 4b) and appears to embay the
spatter cones to the north of the Northwest plateau. Since the elongate plateau, graben, and pits would have been emplaced in the
same eruptive phase if they erupted from the same dike, and the
Northwest plateau would have been emplaced contemporaneously
with the spatter cones for the same reason, this indicates that the
Northwest plateau was emplaced after the outermost drop
K.E. Scanlon et al. / Icarus 237 (2014) 315–339
323
4. After the meltwater drained, the source fissure was no longer
under glaciostatic pressure and, as at the Northwest plateau,
began to erupt locally as a fire fountain, building a tephra
mound and ridge, and eroding a taller melt cavity into the surrounding ice due to more efficient heat transfer in explosive
eruptions.
5. After the eruption ceased, melt from the cavity over the cooling
tephra mound and ridge drained through sinuous englacial tunnels, carrying some tephra and leaving eskers when the glacier
ablated (Fig. 10a and b).
Fig. 7. (a) Topographic profile across the line A–A0 in Fig. 6a. Topographic data from
HRSC-derived DEM. (b) Topographic profile across the line B–B0 in Fig. 6a.
Topographic data from HRSC-derived DEM.
moraines (i.e., those between the bold-outlined ‘‘prominent moraines’’ and the very outermost moraine in Fig. 4b) but before the
bifurcated ridged lava flow, whereas the elongate plateau was
emplaced after the bifurcated ridged flow.
This plateau appears to have formed by a mechanism broadly
similar to that of the Northwest plateau, with a few important differences. We interpret the history of the plateau and its surroundings as follows.
1. Before the emplacement of the elongate plateau, flows associated with the spatter cones to the west of the elongate plateau
and north of the Northwest plateau interacted with remnant ice
in the glacial moraines of the ridged facies. Steam explosions
caused by this interaction imparted a chaotic texture to the
flows. There is the possibility that peperite flows could be
formed in this environment (e.g. Skilling et al., 2002), but orbital
resolution is insufficient to document this. The flows sharply
incised the terminal moraine of the FSD, and continued past
the edge of the deposit.
2. Lava effused from a dike-related fissure below thick glacial ice
cover, forming a steep-sided, elongate plateau enclosed by a
thin melt cavity.
3. Since this vent is much closer to the syneruptive glacier margin
than the vent at the Northwest plateau is, the lava flow near the
edge of the glacier would still have been partially molten when
the melt cavity reached the glacier edge, contacted the atmosphere, and rapidly depressurized. Partial fragmentation of the
flow resulted (Wilson and Head, 2002; Wilson et al., 2013).
Simultaneously, the melt cavity drained catastrophically. The
meltwater entrained the chilled lava fragments and deposited
sediment along a channel downslope of the plateau.
The lack of a well-defined contact between the channel deposits
and the chaotic volcanic flows allows several alternate possibilities
for the history of the region downslope of the elongate plateau
(Fig. 12). First, what we have mapped as channel sedimentary
deposits has been mapped by others as part of the volcanic flow
associated with the cones (Garry et al., 2013). This, however, would
require the flow to have breached several moraines while continuing in a narrow, upslope-directed flow, rather than spreading
between moraines. The ability of such a flow to penetrate beneath
the glacier surface, as indicated by the superposition relationship
observed between the channel and the prominent moraine nearest
the elongate plateau, rather than pooling and spreading laterally at
the glacier edge, is also doubtful. Second, the channel and the flows
may have formed almost simultaneously, allowing mixing and
interaction between the fluvial streams and volcanic flows. This
second interpretation is in conflict with the stratigraphic relationships between the line of cones and the pit craters along the strike
of the elongate plateau (Fig. 4), unless (1) the apparent embayment
relationship between the line of cones and the bifurcated flow is
spurious or (2) there were significant delays between episodes of
eruptive activity at the northern and southern ends of the two
inferred dikes (e.g. Wilson and Head, 1988; Parfitt and Wilson,
1994), such that the cone-associated flows and elongate plateau
were emplaced near-simultaneously but the cones farther north
were emplaced before the pit craters north of the plateau. Our preferred explanation summarized above, i.e. that the flows associated
with the cones were emplaced before the channel associated with
the elongate plateau, is based on these considerations.
3.3. L-shaped ridge
The largest steep-sided flow within the Arsia FSD (Fig. 1a) is a
broad, flat-topped ridge near the northern edge of the deposit
(Fig. 1b, box 3.3). The ridge (Fig. 13), which is primarily L-shaped
in plan view, is 53 km long with a maximum width of 4 km.
Its tallest point lies more than 400 m above the surrounding terrain; with the exception of local highs at the bends in the ridge,
its height decreases gradually downslope from this point
(Figs. 13a and 14). The ridge is much steeper than the two plateaus
at the western edge of the FSD, with some faces inclined at angles
up to 38° (Fig. 13c). The distal branch is leveed (Figs. 13d and 14),
and a small, flat mound extends from the terminus of the ridge. An
approximately triangle-shaped lobe at the upslope end of the ridge
(Fig. 13) has an area of 160 km2, with sides 16 to 28 km long,
and stands 150–200 m tall.
Similar elongate, cohesive, steep-sided ridges have been
mapped in the fan-shaped deposits of Ascraeus Mons and Pavonis
Mons (Zimbelman and Edgett, 1992; Scott et al., 1998); these have
been interpreted to be englacial lava flows (Head and Wilson,
2002; Wilson and Head, 2002; Shean et al., 2005; Kadish et al.,
2008). The Mazama Ridge extending from Mount Rainier (Fiske
et al., 1963; Lanphere and Sisson, 1995; Lescinsky and Sisson,
1998; Lescinsky and Fink, 2000), while composed of more silicic
lava than any known Tharsis flows, is also an instructive analog.
Mazama Ridge is up to 450 m high and up to 1 km wide. It
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K.E. Scanlon et al. / Icarus 237 (2014) 315–339
Fig. 8. (a) The melt produced by heat transfer from lava to ice in a sill-like subglacial eruption (center) is initially confined to a lens above the flow. Reproduced from Fig. 1 in
Wilson and Head (2007). (b) As the lava melts the ice above it, forming an ice cauldron as the thinning ice subsides, a subglacial eruption proceeds from effusive to explosive.
Reproduced from Fig. 10 in Wilson and Head (2002).
formed when lava flowed through a relatively thin corridor of glacial ice: the lava melted ice ahead of it, the melt was able to drain
away, and the lava filled the void. Thicker ice at either side confined the flow, and a glassy, chilled layer formed at its sides and
top while the interior remained molten and continued to flow
downslope. When the flow was temporarily dammed, e.g. by thickness variations in the ice through which it traveled, the lava pooled
and the ridge became locally thicker. Similar terrestrial features in
basaltic eruptions have been described by Wilson and Head (2002).
On the basis of its steep sides, its increased thickness at bends,
and the leveed appearance and thin lobe at the end, we interpret
the L-shaped ridge as a subglacial lava flow. Due to the pinchedoff morphology where the L-shaped ridge and triangular plateau
meet, we propose that the ridge began as a breakout flow from
the plateau and proceeded downslope, possibly following a basal
tunnel eroded by melt from the eruption of the triangular lobe.
The direction of least resistance may also have been set by a crevasse in the ice, a local minimum in the glacier thickness above
the flow, or the slope of the underlying, pre-glacial topography.
The edges of the flow, in contact with glacial ice, would have
chilled much faster than the center.
The morphology of the ridge is consistent with a history in
which, after proceeding downslope (approximately northward)
for 15 km, the flow was temporarily stalled and pooled, resulting
in a local maximum of ridge thickness. Eventually, a breakout flow
emerged and continued approximately westward. After flowing
west for 15 km, the flow stalled and pooled again, creating
another tall and thick region of the ridge, then continued northwest along a sinuous path. As magma supply to the flow ceased,
the still-molten lava in the interior of the ridge partially drained
from the chilled exterior of the distal branch, and the ceiling of
the empty tube would likely have collapsed under the weight of
the glacier, leaving the channels now observed at the distal end
of the ridge.
3.4. Hidden plateaus
Scott and Zimbelman (1995) noted the presence of two mountains buried in the knobby facies, as well as one in a region we have
mapped as belonging to the Smooth Lower Western Flank unit
(Fig. 1b, boxes 3.4). The MOLA and HRSC topographic data that
has since become available revealed a third plateau in the knobby
facies (Fig. 15b). Two of the buried mountains are similar in
morphology to the Northwest plateau, with a small, steep mound
elevated above a flat sheet (Fig. 15a and c). The third lacks a mound
(Fig. 15b).
Any further distinguishing features are obscured by the
superposed knobby facies. In light of the similarities between their
K.E. Scanlon et al. / Icarus 237 (2014) 315–339
325
Fig. 9. Interpretation of the evolution of the Northwest plateau and associated landforms, in plan view (top) and profile view (bottom). (a) The eruption of the Northwest
plateau begins as a pillow lava effusion. The growing pillow sheet embays a crater to its north as it grows, and a lens of liquid water grows around it. (b) As the meltwater lens
grows, it begins to extend under the glacier, allowing the glacier to flow faster locally (as indicated by moraines later left by this part of glacier, shown in gray in plan view
sketch, which bow out from surrounding moraines) and to slide along its bed, streamlining knobs in the direction of flow (brown in plan view sketch). (c) When the meltwater
layer at the base of the glacier reaches the glacier margin, the melt cavity drains catastrophically, forming jokulhlaup channels. (d) After making contact with the atmosphere,
the eruption proceeds explosively under atmospheric pressure, eroding a tall melt cavity in the glacier above the vent. (e) After the eruption ceases, the cooling edifice
continues to melt the surrounding glacial ice. Some of this melt drains in subglacial channels, carrying hyaloclastite tephra from the later stages of the eruption with it. These
channel sediments are left behind as eskers when the glacier sublimes. (For interpretation of the references to color in this figure legend, the reader is referred to the web
version of this article.)
morphology, aspect ratio and size and those of the northwest and
elongate plateaus, however, it is probable that the hidden plateaus
were also emplaced subglacially, by a similar mechanism (Fig. 9). If
so, they would have melted a proportional quantity of ice; the hidden plateaus represent at least 70 km3 of additional subglacially
emplaced lava. The presence of mounds on two of the plateaus
suggests that these eruptions also reached an explosive phase; this
could have been permitted by drainage of that melt at the edge of
the glacier if the edge was close to the plateaus at the time of eruption, or if the eruption melted close enough to the glacier surface to
significantly reduce the pressure above the vent. While no fluvial
features are apparent near the hidden plateaus, channels the size
of those at the northwest and elongate plateaus would easily be
obscured by the superposed knobby facies, and any meltwater that
remained confined by the glacier would have eventually refrozen
in place without leaving any diagnostic landforms. A subaerial lava
cap would be distinguishable in morphology and slope maps even
through the knobby facies, if one had been emplaced (forming a
tuya), but no such feature is apparent on either mound.
3.5. Volcano–ice interactions in the Lobate Facies
The Lobate Facies (Zimbelman and Edgett, 1992) is 7500 km2
in area, located at the northeastern edge of the deposit (Fig. 1b, box
3.5), and consists largely of overlapping sinuous, leveed lava channels and chains of pit craters (Figs. 16 and 17). The unit is named
for lobate, heavily channeled, steep-sided flow-like features that
can be traced tens of kilometers from pit crater chains near the
summit of the volcano (Fig. 16). These features have steep sides
and stand 300 m or more above the surrounding terrain. Aside
from being somewhat thinner, the lobate features at Arsia Mons
are morphologically indistinguishable from similar features at
Pavonis Mons and Ascraeus Mons (Kadish et al., 2008; Shean
et al., 2005).
In light of recent higher-resolution image and topographic data,
the lobate flows at Ascraeus Mons and Pavonis Mons have been
interpreted as subglacially emplaced lava flows (Kadish et al.,
2008; Shean et al., 2005); their stratigraphic position, broad,
steep-sided morphology, and surficial channels are consistent with
this origin, but not with earlier suggestions that they were eskers
(Scott et al., 1998) or pyroclastic flows (Zimbelman and Edgett,
1992). The lack of fluvial features associated with the flows raised
some questions about this hypothesis (Shean et al., 2005); Kadish
et al. (2008) suggested, however, that melt quantities may have
been minimal and that any melt could have locally ponded and
re-froze. We note that terrestrial subglacial basaltic flows can melt
a volume of ice comparable to or greater than their own volume
(Höskuldsson and Sparks, 1997; Wilson and Head, 2002; Tuffen,
2007); although the lower temperature of ice on Mars would
reduce the melt yield somewhat for these flows, the melt production would still have been of the same order of magnitude as the
flow volume (see Section 5). We propose that the melt generated
by these flows, which are far from the downslope FSD margin,
re-froze in place without carving fluvial features.
Many of the steep-sided flows in the Arsia Mons FSD lie along
the strike of pit crater chains farther upslope (Fig. 16). We propose
that this morphology is consistent with the landforms originating
from eruptions beneath the upslope margin of the glacier, with
the transition between pit craters and steep-sided flows marking
the transition from thin ice, firn, or snowpack in the overlying
paleoglacier to thicker, more consolidated ice farther towards the
center of the deposit. The thicker ice farther downslope would provide sufficient confining pressure over the vent to allow an effusive
eruption at the downslope end of the rising dike. The more permeable snow, ice, or firn farther upslope would allow meltwater to
interact with the rising magma at near-atmospheric pressure,
causing explosive phreatomagmatic eruptions to form a chain of
pit craters (Wilson and Head, 2007a,d).
Near the downslope extent of the Lobate Facies lies a steepsided plateau, 13 km long and 3 km wide, tapering in height
from 300 m above the surrounding terrain at its upslope extent
to 100 m above the surrounding terrain at its terminus. Near
the terminus, the center of the plateau has a sunken, leveed
appearance, and a thin, delta-shaped lobe 4 km long by 5 km
wide emerges at the distal end (Fig. 17). Based on its flat top and
steep sides upslope, we suggest that this ridge emerged as a subglacial ice-confined lava flow similar to the L-shaped plateau. Like
the L-shaped plateau, this feature has a leveed, sunken center near
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K.E. Scanlon et al. / Icarus 237 (2014) 315–339
Fig. 10. The elongate plateau. (a) HRSC DEM superimposed on CTX images. (b) Sketch map showing location of topographic profiles in Fig. 11. (c) Slope map, derived from an
HRSC DEM. Contacts between the ridged and knobby facies are shown in white; other geological features are outlined in black.
K.E. Scanlon et al. / Icarus 237 (2014) 315–339
Fig. 11. (a) Topographic profile across the line A–A0 in Fig. 10b. (b) Topographic
profile across the line B–B0 in Fig. 10b. (c) Topographic profile across the line C–C0 in
Fig. 10b. Topographic data from HRSC-derived DEM.
its terminus. We interpret this as resulting from part of the molten
lava core draining downslope. The delta-shaped distal lobe could
have formed from this late stage subglacial breakout and drainage
of the steeply leveed flow, filling accommodation space provided
by the melt cavity in the glacier.
3.6. The Smooth Lower Western Flank unit
The region mapped by Zimbelman and Edgett (1992) as the
Smooth Lower Western Flank unit (Fig. 1a, yellow) is generally distinguished by flat lava deposits similar to those erupted subaerially
(Fig. 1b, box 3.6). It also contains several steep-sided structures
with morphologies not seen elsewhere in this deposit. It borders
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Fig. 12. Channel and putative hyperconcentrated flow deposits emerging from the
elongate plateau and embaying a line of spatter cones to the northwest of the
plateau. (a) CTX image mosaic. Prominent moraines cross-cut by the channeled
debris are marked with black arrows; the terminal moraine cross-cut by the
hummocky flow is marked with a white arrow. (b) Sketch map.
the Degraded Flank unit (Section 3.7), which is characterized by
a series of ‘‘irregular scarps’’ (Zimbelman and Edgett, 1992).
One such feature (Fig. 18) is an exceptionally smooth-topped
plateau, 9 km long by 5 km wide and 150 m high (Fig. 18b
and c), with sides inclined at angles 12–23° where HRSC DTM
coverage is available (Fig. 18d). Three small craters (200–500 m
in diameter) lie along the center of the approximately equidimensional plateau; the plateau surface is otherwise as smooth as the
surrounding, apparently subaerial, flows. This feature resembles a
terrestrial subglacial volcanic landform known as a tuya, or table
mountain.
Tuyas (e.g. Gudmundsson, 2003; Jones, 1968; Smellie, 2007) are
landforms that begin as pillow mounds topped by hyaloclastite
tephra, with eruptions proceeding in a manner similar to the
Northwest, Elongate, and buried plateaus in the western part of
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Fig. 13. The L-shaped ridge and surrounding landforms. (a) THEMIS daytime IR image mosaic. (b) Sketch map of the L-shaped ridge. (c) HRSC DEM-derived slope map of the
portion of the ridge with HRSC DEM coverage. (d) CTX image mosaic detail view of the L-shaped ridge.
the Arsia FSD (Fig. 1). If the eruption lasts long enough that the
melt cavity penetrates to the surface of the ice sheet, the eruption
will continue subaerially. A cohesive capping flow will form atop
the tephra mound, and deltas of quenched lava will grow outward
where the capping flow enters the water in the melt cavity, now an
ice-confined subaerial lake. We interpret the exceptionally smooth
plateau as a tuya. The shape of the feature in profile (Fig. 18c)
strongly resembles those of mafic tuyas in Iceland, British Columbia and Antarctica, such as the Herðubreið tuya (e.g. Werner and
Schminke, 1999), Tuya Butte (e.g. Mathews, 1947), and the Jonassen tuya (e.g. Smellie, 2006), respectively. The height of the plateau
is somewhat less than those of terrestrial mafic tuyas with the
same area (Fig. 19; Smellie, 2009). This could indicate that the
explosive phase of the eruption was short, leaving a landform similar to a pillow mound with a subaerial cap; that the ice cover was
thin while this plateau was built, allowing a subaerial phase to be
reached without much vertical edifice construction; or that the iceconfined lake level, which determines the height of the lava deltas,
was lower than in terrestrial settings. The largest ‘‘crater’’ on the
surface of the plateau resembles a vent in the slope map
(Fig. 18d). Since the plateau is so close to the summit of Arsia Mons
(200 km from the caldera center), the glacial ice covering it during the Arsia tropical mountain glacier period would have been
comparatively thin, less than 1.8 km even if the eruption
occurred with the glacier at its maximum extent (Shean et al.,
2005; Fastook et al., 2008). Depending on the time of emplacement, the comparatively thin ice in this region may have allowed
the eruption to reach an explosive stage rapidly; the tephra would
have melted the ice above it rapidly relative to effusive eruptions
in deeper ice, quickly breaching the glacier surface and enabling
the capping flow to form subaerially.
Another morphology common in the Smooth Lower Western
Flank unit is that of the sunken-centered, steep-sided flow. These
features resemble the L-shaped and smooth-topped plateaus in
planform, height, and steepness, but are much lower in their interiors than at the margins, sometimes level with the surrounding
K.E. Scanlon et al. / Icarus 237 (2014) 315–339
Fig. 14. Topographic profiles of the L-shaped ridge. (a) Topographic profile across
the line A–A0 in Fig. 13b. (b) Topographic profile across the line B–B0 in Fig. 13b. (c)
Detrended topographic profile across the line C–C0 in Fig. 13b. Topographic data
from HRSC-derived DEM.
plains (Fig. 20a–c). Only one sunken-centered flow lies outside the
Smooth Lower Western Flank, in the Lobate Facies. We interpret
this general morphology to result from lava flow emplacement,
with steep chilled margins of the flow forming during flow
emplacement, growth and inflation. The decrease in topography
in the flow interior could be due to a variety of factors, including
subsequent phases of flow emplacement, breakouts, distal drainage (note the rubbly terrain surrounding several of these features),
and lava withdrawal back into the vent.
Several low mounds (Fig. 21a and b), like those forming the base
of the glaciovolcanic plateaus on the western edge of the FSD, are
also present in the Smooth Lower Western Flank (Fig. 1a; yellow).
The regional slope in this unit is significantly steeper than at the
west of the FSD, and nearly all of the mounds are elongated in
the downslope direction. Mound maximum heights range from
100 to 200 m; some slope gently, and others have steep edges
329
comparable to those of the plateaus to the west. Some mounds
are crosscut by local faults and others superimpose the faults
(Fig. 21a and b). Due to their size, morphology, and morphometry,
we interpret these features as sill-like pillow sheets and pillow
mounds (e.g. Smellie, 2007; Cousins et al., 2013), the products of
local subglacial eruptions.
One feature, located at the contact between the Smooth Lower
Western Flank and Knobby Facies and partially obscured by
Knobby Facies material (Fig. 22a and b), is much larger than the
others: 86 km long by 8–30 km wide and 200 m high, with
an approximately circular superimposed mound about 16 km2 in
area and 100 m high. The long axis of the feature is approximately
parallel to those of the Large Graben to the east and Aganippe Fossa
to the west. Where HRSC DEMs are available, they show its edges
to slope as steeply as 15°. One of the sunken-centered flows
(Fig. 20b) lies on its southwestern corner. The thickness of this flow
and the steepness of its edges surpass those of subaerial flows, but
are comparable to the hypothesized pillow effusions found
throughout this region of the deposit. We therefore interpret the
feature as a mound of pillow lava erupted from a subglacial fissure.
East of this feature, also spanning the boundary between the
Smooth Lower Western Flank and the Knobby Facies, lies a partially buried graben, similar in appearance to the Large and Small
Graben and Aganippe Fossa (Fig. 22). The graben trends southwest-to-northeast, parallel to the Small Graben but in contrast to
the predominantly north-to-south trend of other, similarly-sized
graben in the deposit. West of this partially buried graben lies a
steep, irregular scarp 85 km long. At its southern end, a curved
section of the scarp stands up to 400 m tall. We interpret this scarp
as the edge of a subglacial flow predating the partially buried graben. A pair of sunken-centered flows superpose the graben and the
material burying it.
Three plateaus in this region are associated with these graben
(Fig. 1b). Two extend from the Large Graben, near the contact
between the Degraded Flank and Smooth Lower Western Flank
units (Fig. 23a). The larger of these two plateaus extends 20 km
upslope from the east side of the Large Graben. It is 9 km wide,
up to 600 m high, and steep sided, with slopes of 20–30° common
on all faces covered by HRSC DEMs. It is incised by a deep, sinuous
channel at least 8 km long; discontinuous ridges and hollows elsewhere may be sections of the same, partially buried, channel. The
taller part of the plateau is smooth but covered in small dunes;
the lower part has a hummocky appearance. Below the plateau
and level with its base, the graben walls expose thin, horizontal
layered flows with relatively even spacing (Fig. 23b and c). In
CTX images stretched to enhance contrast at the edge of the plateau, the graben exposes several cohesive lava flows, but talus is
abundant, especially higher in the section. Several layers within
the talus are brighter than those above and below (Fig. 23b). The
bright layers may correspond to compositional or grain size variations in the unconsolidated material. A few of these layers appear
to be associated with outcrops of cohesive material.
Based on the morphology of the plateau and its abundance of
unconsolidated material relative to the stratigraphically lower
flows exposed by the graben (Fig. 23b), we propose that the plateau was built by explosive subglacial eruptions. If this is the case,
the tephra is more abundant higher in the sequence because the
confining glacial pressure suppressed explosive activity earlier in
the eruption, as we interpret to have occurred in the plateaus to
the west. We predict that further inspection of this plateau with
higher-resolution images will reveal that the consolidated layers
in the plateau are substantially thicker than those below it, indicating that they were emplaced in an ice-confined vault, and that they
may contain pillow structures.
The smaller, knobby mound (Fig. 23a and d), extending 4 km
outwards from the south end of the Large Graben, is 10 km wide
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K.E. Scanlon et al. / Icarus 237 (2014) 315–339
Fig. 15. Plateaus buried in the knobby facies (left) and detrended profiles across the plateaus (right). The plateaus in (a and b) have steep mounds morphologically similar to
those on the Northwest and elongate plateaus. MOLA elevation map superimposed on CTX images; profile data from 128° per pixel MOLA gridded dataset.
K.E. Scanlon et al. / Icarus 237 (2014) 315–339
331
Fig. 17. Steep-sided plateau in the Lobate Facies. This plateau follows the same
downslope (south to north) progression from a flat-topped ridge, to a sunkencentered ridge, to a thin distal lobe as the L-shaped ridge. (a) CTX image mosaic. (b)
Sketch map.
Fig. 16. Steep-sided, lobate flows grading into pit crater chains in the lobate facies.
(a) CTX image mosaic. (b) Sketch map; steep sided flows are shown in black, thinner
lava flows are shown in gray, pit craters are shown in blue, and lava channels are
shown in red. (For interpretation of the references to color in this figure legend, the
reader is referred to the web version of this article.)
and 150–200 m tall. Its surface texture is knobby, and its edges
are not as steep as those of the larger plateau to the east. Thermodynamic relationships (Wilson and Head, 2007a) and observations
of recent terrestrial eruptions (Belousov et al., 2011; Edwards et al.,
2012) suggest that supraglacial lava flows can advance kilometers
before melting through the snow or ice beneath them. If the flow is
able to solidify before the snow or ice is completely removed, the
flow will fracture into large blocks as it collapses (Edwards et al.,
2012).
Based on this terrestrial analogue and the occurrence of a similar knobby texture at the edges of hypothesized ice-marginal
flows elsewhere in the deposit (Section 3.7), our preferred explanation for the knobby mound south of the Large Graben is that it
formed as a supraglacial effusive lava flow from a vent within
the graben, and gained its current appearance when the ice
beneath it was completely removed. The gently sloping sides of
this mound could have resulted if the flow was thinner at the edges
due to partial melting into the glacier during emplacement, or from
mass wasting of loose rubble after the ice was removed.
Alternatively, the knobby mound could have been formed by
the emplacement of volcanic tephra from an explosive eruption
in the graben. Aganippe Fossa and the other morphologically similar graben in the Arsia Mons FSD, including the Large Graben, may
have been formed by large-scale phreatomagmatic eruptions
caused by the interaction between the rising magma and meltwater generated by dike emplacement into the glacier (Wilson and
Head, 2007d). Hot tephra would have melted the ice below it
(Wilson and Head, 2007e), potentially producing the observed
uneven surface texture. If this was the formative mechanism for
the knobby mound, however, the plan-view shape of the feature
and the lack of similar features near other graben in the deposit
are unexpected.
The Smooth Lower Western Flank unit (Fig. 1a, yellow) is generally distinguished by a lack of the glacial debris that forms the
knobby facies and ridged facies at lower altitudes in the deposit.
In this upslope region of the Late Amazonian Arsia Mons glacier,
climate and glacial models indicate that snow accumulation
exceeded sublimation for most of the glacier’s life span (Fastook
et al., 2008), whereas sublimation exceeded accumulation in the
downslope regions (ablation zone) now occupied by the ridged
and knobby facies. Mouginis-Mark and Rowland (2008) studied
High Resolution Imaging Science Experiment (HiRISE) imagery of
the Small Graben in the Smooth Lower Western Flank unit
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K.E. Scanlon et al. / Icarus 237 (2014) 315–339
Fig. 18. (a) Flat-topped plateau in the Smooth Lower Western Flank unit. The distinct boundary between the capping unit and the unit forming the side slopes is highlighted
with white arrows. CTX image. (b) CTX image of the plateau in (a) shaded by HRSC-derived elevation. (c) Topographic profile across the line A–A0 in (b). Topographic data from
HRSC-derived DEM. (d) Slope map of the plateau derived from HRSC DEM data shown in (b).
(Fig. 1b) and noted that no layers among the sequence of lava flows
exposed in the graben wall were clearly glacial in origin. While
they noted the presence of an unconformable layer that could be
glacial sediment, as well as the possibility that the entire sequence
exposed by the graben could have been emplaced between periods
of high obliquity, we suggest that the lack of obvious glacial debris
in the cross-section exposed by the Small Graben is a function of its
location in this upslope region of the glacier, where little glacial
debris is visible even at the surface.
3.7. Digitate cliffs
Except in the Lobate Facies, the upslope edge of the FSD (Fig. 1b,
box 3.7) is marked by numerous overlapping, digitate, flat-topped,
steep-sided protrusions (Fig. 24). A typical protrusion is a few kilometers wide and tens of kilometers long. Side slope angles, where
HRSC DEM coverage is available, are as steep as 20°, though 10°
slopes are more common. The protrusions farthest downslope are
the lowest stratigraphically, and they appear subdued relative to
the younger ones. The stratigraphically highest protrusions are
superposed only by the youngest subaerial flows, pit craters, and
concentric graben at the summit of Arsia Mons (Fig. 24a and b).
The margins of some protrusions are similar in texture to the
knobby mound, and in some cases these knobby margins appear
to superpose graben that crosscut the surrounding terrain and
the interior regions of the same flows.
Similar digitate features appear at the summits of Ascraeus
Mons and Pavonis Mons, but only on the sides of their summits
in contact with their fan-shaped deposits (Kadish et al., 2008;
Fig. 24c). Previous workers have interpreted these cliffs as erosive
features caused by a mass flow or high-temperature lavas
(Zimbelman and Edgett, 1992), or as analogous to the Olympus
Mons basal scarp (McGovern and Solomon, 1993), though Scott
and Zimbelman (1995) noted that they more nearly resembled
individual flows emplaced prior to the flows above them. On Earth,
the Hoodoo Mountain composite volcano in British Columbia is
ringed by cliffs up to 200 m tall, whose steep edifices, glassy surface texture, and columnar jointing indicate that they formed as
K.E. Scanlon et al. / Icarus 237 (2014) 315–339
333
Fig. 19. The volume-to-area ratio of the smooth plateau (red triangle) lies below
the trend of terrestrial tuyas (black diamonds) and is more similar to pillow mounds
and ‘‘sheet-like sequences’’ (black circles and squares). Reproduced and annotated
from Smellie (2009). (For interpretation of the references to color in this figure
legend, the reader is referred to the web version of this article.)
subaerial summit lava flows flowed downslope and chilled against
a glacier that surrounded the mountain (Edwards and Russell,
2002; Edwards et al., 2002).
On the basis of their morphology and their exclusive association
with the contact between subaerial summit lava flows and the
edge of the FSD, we interpret the Tharsis Montes digitate cliffs as
ice-margin chilled lava flows analogous to those at Hoodoo Mountain (Fig. 25). While the Hoodoo Mountain lavas, like the Mazama
Ridge lavas we compare to the L-shaped ridge above, are more
silicic than we expect Tharsis lavas to be, basalt flows also develop
chilled crusts where they contact and melt ice, confining the flows
and producing steep walls of pillow lavas (e.g. Edwards et al., 2009;
Skilling, 2009). Higher-resolution images than are presently
available for the Arsia Mons cliffs may reveal pillow structures or
jointing that will allow more detailed investigation of the lava
composition and cooling history; higher resolution DEMs will also
provide more insight into the structure of these cliffs. We propose
that the knobby margins of the cliffs are regions where the flow
overrode the glacial ice, leaving rubble piles when the glacier
receded. The superposition relationships of these margins can be
explained if the tectonic events creating the graben occurred while
ice was still present: when the ice eventually sublimated, the flows
that had been supported by ice would have fractured, burying the
segments of graben beneath them in rubble.
The digitate cliffs at Arsia form three discrete tiers, with the tier
farthest downslope forming the contact between the Degraded
Flank and Smooth Lower Western Flank units as we have mapped
them (Fig. 1) and the tier farthest upslope forming the contact
between the summit and Degraded Flank units; the middle tier
and several discontinuous cliffs fall within the Degraded Flank
unit. In terrestrial glaciers, ice can accumulate only above the equilibrium line altitude (ELA), the altitude where accumulation of
snow into the glacier is equal to the ablation of the glacier by melting or sublimation (Benn and Evans, 2010). On Mars, the increase
of sublimation with altitude is often greater than the decrease of
temperature with altitude, leading to a second, higher-altitude
ELA above which ice does not accumulate (Fastook et al., 2008).
Fig. 20. Sunken-centered, steep-sided flows in the Smooth Lower Western Flank
unit (a and b) and the Lobate Facies (c). HRSC DEM (where available) superimposed
on CTX images.
334
K.E. Scanlon et al. / Icarus 237 (2014) 315–339
Fig. 21. Steep-sided and gently sloping mounds in the Smooth Lower Western
Flank unit. (a) MOLA elevation data superimposed on a CTX image mosaic. (b)
Sketch map.
We infer that each tier of digitate cliffs represents a series of Arsia
Mons summit eruptions in a climate with a different upper ELA for
the glacier. The 4 km difference between the altitude of the upslope limit of the glacier when the highest and lowest tier were
emplaced implies that the balance between temperature, atmospheric pressure, and precipitation changed significantly between
the first and third summit eruptive episodes represented by the
cliffs (Fig. 25), though the influences of these climate variables cannot be deconvolved a priori (Fastook et al., 2008). The wide separation between each of the three layers is in keeping with other
evidence that construction of the Tharsis Montes was episodic
(Wilson et al., 2001).
Fig. 22. Feature interpreted as a large pillow mound and associated landforms in
the Smooth Lower Western Flank unit. (a) THEMIS daytime IR image mosaic. (b)
Sketch map; impact craters are shown with triple lines, drop moraines left by latestage graben glaciers are shown with thin black lines, volcanic graben are shown
with thick red lines, tectonic graben are shown with thick blue lines, geologic
contacts are shown with thick white lines, and other geologic features are shown
with thick black lines. (For interpretation of the references to color in this figure
legend, the reader is referred to the web version of this article.)
flow of the overlying ice (cold-based to wet-based flow), and the
meltwater can be released catastrophically in jökulhlaups. The
quantity of ice melted by the subglacial eruptions recorded in
the Arsia Mons fan-shaped deposit, and the longevity of the resulting liquid water bodies, are of primary astrobiological interest. We
use thermodynamic relationships, validated by terrestrial measurements, to estimate the volume of meltwater produced in the
construction of several Arsia Mons subglacial edifices, and the time
scale over which this water remains liquid (e.g., Wilson and Head,
2002, 2007a, 2009; Gudmundsson, 2003; Head and Wilson, 2007;
Tuffen, 2007; Wilson et al., 2013).
The heat energy E released by the rapid chilling of a volume Vm
of magma (Gudmundsson, 2003) is given by:
4. Meltwater: volume and longevity
E ¼ q m V m c g DT m ;
Interaction of ascending and intruding magma with glaciers and
glacial ice deposits produces significant quantities of meltwater
that can be stored and transported subglacially, influencing the
where Vm is the volume of magma, qm is the density of magma, cg is
the specific heat capacity of the glass formed by rapid chilling of the
magma, and DTm is the temperature change the magma undergoes.
ð1Þ
K.E. Scanlon et al. / Icarus 237 (2014) 315–339
335
Fig. 23. (a) The Large Graben. The large, eastern plateau is denoted with a black arrow, and the knobby mound is denoted with a red arrow. CTX image mosaic. (b) The plateau
east of the Large Graben, as exposed by the Large Graben. The unconsolidated material is much thicker in the vicinity of the plateau than in the adjoining plains (white arrow).
Brightness variations in the layers of the plateau are highlighted with a red arrow. CTX image, stretched to enhance contrast in bright pixels. (c) Sketch map of the region
depicted in (b). The plateau is shown in blue, the edge of the graben in purple, the edge of the graben glacial deposits in green, thinly layered flows in red, cohesive outcrops in
black, and bright layers in gray. (d) The knobby mound south of the Large Graben. CTX image mosaic. (For interpretation of the references to color in this figure legend, the
reader is referred to the web version of this article.)
We use the specific heat of basaltic magma as a conservative
estimate of the specific heat of basaltic glass (Gudmundsson,
2003). The heat energy Q required to melt a volume Vi of ice
(Gudmundsson, 2003) is given by:
Q ¼ qi V i ½ci ðT 0 T 1 Þ þ Li þ cw ðT 2 T 0 Þ;
ð2Þ
where qi is the density of ice, ci is the specific heat capacity of ice, T0
is the melting temperature of ice, T1 is the initial temperature of the
ice, Li is the latent heat of fusion of ice, cw is the specific heat capacity of liquid water, and T2 is the temperature of the resulting
meltwater.
The efficiency with which the magma can transfer its heat to
the ice, fmi, is:
336
K.E. Scanlon et al. / Icarus 237 (2014) 315–339
Fig. 24. Steep-sided, digitate flows at the upslope edge of the Arsia Mons fan-shaped deposit. (a) THEMIS daytime IR image mosaic. (b) Sketch map. Steep-sided flows are
outlined in black, pit craters are outlined in violet, impact craters are outlined in blue, tectonic graben are outlined in green, and the Small Graben is outlined in red. Knobby
flow margins are shaded in gray, and generally appear to superpose tectonic features that cross-cut the steep-sided flows. (c) Comparison between digitate flows at the
upslope edge of the Arsia Mons (left), Pavonis Mons (center) and Ascraeus Mons (right) fan-shaped deposits, plotted at the same scale for comparison. THEMIS daytime IR
mosaics. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)
337
K.E. Scanlon et al. / Icarus 237 (2014) 315–339
We make the following assumptions: (1) the thermodynamic
properties of lavas in the Arsia FSD were similar to those of terrestrial basalts (Table 1), (2) the internal temperature of the ice was
similar to the modern martian equatorial surface temperature at
high obliquity, (3) integrated over the course of the eruption,
fmi Q/E, and (4) the final temperature reached by the water is
15 °C, typical of terrestrial eruptions (Gudmundsson, 2003;
Tuffen, 2007). Then the volume of ice melted by a volume of
magma Vm is equal to:
Vi ¼
fmi qm V m cg DT m
;
qi ½ci ðT 0 T 1 Þ þ Li þ cw ðT 2 T 0 Þ
ð4Þ
for a corresponding volume Vw of meltwater:
Vw ¼
fmi qm V m cg DT m
;
qw ½ci ðT 0 T 1 Þ þ Li þ cw ðT 2 T 0 Þ
ð5Þ
where qw is the density of water. Under these assumptions, Eq. (5)
indicates that the formation of the Northwest plateau melted
39–46 km3 of ice, producing 36–42 km3 of liquid water: 22 km3
during the formation of the pillow sheet, and 14–20 km3 during
the formation of the central tephra mound. The formation of the
L-shaped ridge, for example, would have melted another 38 km3.
Due to the channel-forming jökulhlaup that occurred at the
Northwest plateau, much of the water formed in the first phase
of that eruption would have frozen or evaporated before the eruption ceased. Some of the water formed in the explosive phase is
also likely to have escaped beneath the glacier before the glacier
refroze to its bed. If most of the water from the explosive phase
remained confined, it may have survived long enough to become
inhabited by any extant surface life.
If, as an example, we model the melt cavity formed by the central mound of the Northwest plateau as a uniform planar sheet
covering the same area as the mound, and use the more conservative estimate of meltwater generated by this phase of the eruption,
the meltwater layer is 500 m thick. The fact that the 150 m tall
central mound did not form a subaerial lava cap suggests that the
meltwater layer was at least 150 m thick. Though lakes should
become increasingly saline as they freeze, we will conservatively
assume fresh water for these calculations. Heat loss from the water
to the cooling edifice below it is likely to be small compared to loss
to the ice above (Wilson and Head, 2002), since conductive heat
flux across a surface is proportional to the temperature gradient
across that surface. Therefore, we consider only heat transfer from
Table 1
Values used in ice melt calculations.
Value
Fig. 25. Schematic depiction of the formation of the Tharsis Montes cliff tiers in a
changing climate. At the upslope edge of the FSD-forming glacier (a), lava flowing
downslope from the summit will form steep cliffs where it chills rapidly against the
glacier (b). If changes in the martian climate (e.g., relative humidity, total
atmospheric pressure) cause the upper equilibrium line altitude (ELA) to rise, later
summit eruptions will form chilled cliffs at a higher altitude on the volcano (c). This
cycle will repeat, forming multiple tiers of chilled flows (d), which remain as cliffs
after the glacier sublimates (e).
fmi ¼
dQ
dt
dE
;
dt
Northwest plateau, pillow sheet
Vm
fmi
3.1 1010 m3
0.1
Northwest plateau, tephra mound
Vm
fmi
3.9 109 m3
0.5–0.7
L-shaped ridge
Vm
fmi
5.2 1010 m3
0.1
Basaltic lava
cg
qm
Tm
ð3Þ
Water
Li
qi
and has been empirically estimated to equal 0.1 for pillow lavas
and 0.5–0.7 for hyaloclastite tephra (Gudmundsson, 2003;
Tuffen, 2007).
Ti
ci
qw
>1200 J kg1 K1
2300 kg m3
1200 °C
3.35 105 J kg1
917 kg m3
60 °C
2050 J kg1
1000 kg m3
338
K.E. Scanlon et al. / Icarus 237 (2014) 315–339
the water to the ice above it. Under these simplifying assumptions,
the time to freeze the melt layer is given by Kreslavsky and Head
(2002):
2
tðhÞ ¼
Li qi h
;
2ji ðT 0 T s Þ
ð6Þ
where Li and qi are the latent heat of fusion and density of ice; h is
the thickness of the melt layer; ji is the thermal conductivity of ice,
2.5 W m1 K1; T0 is the melting temperature of the melt water,
0 °C for fresh water; and Ts is the surface temperature, 60 °C for
Amazonian equatorial Mars at high obliquity. Thus we estimate that
a 150 m thick layer would not freeze through for 730 years, or
8000 years for a 500 m thick layer. Even the lower estimate would
represent a sufficient time for many generations of microbes to
have flourished in this melt body, if they existed to colonize it.
5. Conclusions
We conclude that glaciovolcanic landforms are abundant in the
Arsia Mons fan-shaped deposit. These include landforms interpreted as subglacial pillow sheets larger than any known on Earth,
hyaloclastite mounds, large englacial and glacier-marginal ice-confined flows, and a tuya. The presence of three well-separated strata
of ice-marginal lava flows at the upslope margin of the deposit
indicates that near-summit eruptions occurred in three distinct
sets of climate conditions.
Glaciovolcanism in the deposit would have melted massive volumes of ice at the upslope margin, throughout much of the Lobate
Facies, and at numerous eruptive centers throughout the rest of the
deposit. The best-exposed of these eruptions, at the Northwest plateau, would have created an englacial lake that may have persisted
for hundreds to thousands of years. This meltwater production and
glaciovolcanic heat transfer caused local wet-based flow in the
Arsia Mons glacier (Scanlon et al., in preparation, 2014), creating
additional habitable environments in the Arsia Mons FSD.
Acknowledgments
We gratefully acknowledge support from the NASA Graduate
Student Researchers Program to KES and from the Mars Data Analysis Program and the Mars Express High-Resolution Stereo Camera
Investigation Team to JWH. We thank fellow attendees of the third
IAVCEI-IACS Volcano–Ice Interactions Conference for productive
discussions. We thank M. G. Chapman and an anonymous reviewer
for constructive reviews that improved the manuscript.
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