Icarus 180 (2006) 75–87
www.elsevier.com/locate/icarus
Detailed study of an hydrological system of valleys, a delta and lakes
in the Southwest Thaumasia region, Mars
Nicolas Mangold ∗ , Véronique Ansan
Lab. IDES, Bat. 509, Université Paris-Sud et CNRS, 91405 Orsay, France
Received 20 December 2004; revised 26 April 2005
Available online 9 November 2005
Abstract
The occurrence of fluvial activity and standing bodies of water on early Mars is the subject of debate. Using MOC, MOLA, and THEMIS data,
we identify a whole set of landforms in the Thaumasia region which attest to water flows during geologically long periods of more than thousand
years. A thick fan-delta is identified within an impact crater at the outlet of a deep valley. Ponded water filled and overflowed this crater’s rim,
creating entrance and exit breaches and an outlet valley. These landforms show that the 25-km diameter impact crater contained a lake up to 600
m deep. At the head of this crater’s deep contributing valley, a closed depression may have contained another lake, but depositional landforms
are not evident in this headward basin. Alternatively, groundwater discharge may have supplied the valley, but the observed landforms are not
consistent with a sudden release of water, as is usually invoked for the large martian outflows channels. Stratigraphic relationships show that this
hydrological activity occurred during the Hesperian period, thus relatively late in the history of martian valley network development.
 2005 Elsevier Inc. All rights reserved.
Keywords: Mars, surface; Mars, climate
1. Introduction
The climate of the early Mars has been a subject of debates
for 30 years centered around two end-member models theories
involving a warm and wet climate or cold and dry processes
only (e.g., Pollack et al., 1987). Groundwater flows driven by
early high geothermal gradient are often invoked to explain valley networks in cases where paleoclimate models do not reach
warm temperatures (Kasting, 1991; Clifford, 1993). Nevertheless, recent climate models (Forget and Pierrehumbert, 1997;
Mischna et al., 2000; Colaprete and Toon, 2000) and recent observations of valley networks (Craddock and Howard, 2002;
Grant, 2000; Mangold et al., 2004) or potential paleolakes
(Cabrol and Grin, 1999; Ori et al., 2000a; Malin and Edgett,
2003) renew the case for a warmer early Mars. Potential paleolakes were identified inside impact craters on the basis of
geomorphic criteria and/or topographic data (de Hon, 1992;
Newson et al., 1996; Cabrol and Grin, 1999, 2001, 2002; Irwin
* Corresponding author. Fax: +33 1 69 15 63 48.
E-mail address: [email protected] (N. Mangold).
0019-1035/$ – see front matter  2005 Elsevier Inc. All rights reserved.
doi:10.1016/j.icarus.2005.08.017
et al., 2002) but the longevity of these putative standing bodies of water remains uncertain. A recent study shows the most
convincing evidence for a potential paleolake in the “north–east
Holden crater” which contains a distributary fan or delta with
meandering inverted channels (Malin and Edgett, 2003). Although valley network development occurred primarily during
the Noachian period perhaps associated with warmer climate
(Craddock and Maxwell, 1990), some paleolakes could have
existed during the Hesperian and Amazonian periods (e.g., Ori
et al., 2000a).
In this study, we identify two potential paleolakes using
topographic data of Mars Global Surveyor (MGS) Mars Observer Laser Altimeter (MOLA, Smith et al., 1999), wide angle and narrow angle visible Mars Observer Camera images
(MOC, Malin et al., 1998) acquired at different spatial resolution (∼250 m and 3 m/pixel) and thermal imagery acquired
with the Mars Odyssey Thermal Emission Imaging System instrument (THEMIS, Christensen et al., 2003). Maps, volume
and slope measurements are done using a Geographic Information System (GIS) software after referencing all data in a
martian standard. The two potential paleolakes are located in
the southern part of the Tharsis region inside high-standing ter-
76
N. Mangold, V. Ansan / Icarus 180 (2006) 75–87
rains dated to the Late Noachian epoch. The two basins are
connected by a deep valley which shows a 600-m thick delta
at its mouth (Ansan and Mangold, 2005). Chronological relationships show that this fluvial activity postdates terrains dated
of the Noachian–Hesperian boundary indicating Hesperian age
water flows (Dohm et al., 2001), a period thus relatively late in
the history of martian valley network development.
2. Tectonic and fluvial landforms in West Thaumasia
highlands, South Claritas Fossae region
2.1. Geologic context and tectonic features
Claritas Fossae and Thaumasia highlands are located 39◦ S
and 103◦ W at the southern end of the Tharsis bulge (Fig. 1).
This region consists of heavily cratered Noachian highlands
that were uplifted during the Noachian period, which ended 3.8
to 3.6 Ga ago (Dohm and Tanaka, 1999) and surrounded by
smooth Hesperian plains along their western edge (Dohm et al.,
2001). This uplift has been attributed to outward-verging fold
and thrust margins (Schultz and Tanaka, 1994) or magmatic activity and crustal underplating (Dohm and Tanaka, 1999). The
most abundant geomorphic features over these rugged highlands consist of (1) extensional tectonic faults mainly oriented
in the north–south direction and (2) valley networks such as
Warrego Valles located to the southeast of the studied area
(Dohm et al., 2001; Gulick, 2001; Ansan and Mangold, 2003,
2005). These two kinds of geologic features are well expressed
in the mosaic of daytime THEMIS images, in which the contrast is due primarily to albedo and slope (Fig. 2a). The similarity in tone on the visible image (Fig. 1) indicates a predominance of topographic effect in the THEMIS image. In the
nighttime THEMIS mosaic (Fig. 2b), spatial differences in the
thermal inertia of material are evident but do not organize into
laterally extensive specific geologic units. Bright areas in the
image represent the rocky slopes of faults or craters and other
contrasts may indicate regional differences in the properties of
aeolian mantles.
Well-developed tectonic faults are evident in all context images. The predominant north–south direction is composed of
normal faults forming narrow grabens of a few kilometers in
width and up to 100 km long (Fig. 1). West of the studied area,
some of these narrow grabens (Fig. 1) continue into the smooth
plains southwest of the highland boundary. The part of these
plains visible in Fig. 1 is dated to the Noachian–Hesperian transition (HNf on the map of Dohm et al., 2001). The extensional
tectonic activity, partially cutting these terrains, therefore began
in the Noachian period and continued in the Hesperian epoch
or later. Minor east–west trending normal faults or fractures exist locally in the southern and western parts of the study area.
These grabens are relatively shallow and more degraded, which
may indicate an older age. In the MOLA digital elevation model
(DEM), single grabens typically have depths of less than 100 m,
but a prominent graben oriented NNE–SSW displays a sharp
break in slope with 500 m of difference of elevation (Fig. 3a).
The set of normal faults bounds a large N–S trending closed depression (∼ 250 × 80 km) in the central part of the studied area
(Fig. 3b). This depression is similar to basins observed in rift
zones on Earth (e.g., Needham et al., 1976).
2.2. Fluvial features
Branching valleys similar to terrestrial fluvial networks
are locally observed mainly on the southern slope of the
Noachian period Thaumasia highlands bounding the smooth
plain (Fig. 3b). Valleys show a subparallel drainage pattern
rather than a tree-like pattern, indicating control by the topographic slope (Schumm et al., 1987). THEMIS images show
that fluvial valleys are incised less deeply than are the grabens.
The topography of fluvial valleys is usually not visible at the
MOLA resolution (463 m/cell). Several valleys identified on
the nighttime THEMIS images converge into the eastern closed
depression although these valleys are sparse (Fig. 3b). Valleys could have developed in two or more stages from the
Noachian into the Hesperian periods (Dohm and Tanaka, 1999).
Such valleys were first attributed to rainfall-fed fluvial erosion (Sagan et al., 1973). Hypotheses involving hydrothermal
activity associated with impacts, volcanism or tectonics have
been proposed to explain morphological differences between
martian valleys and terrestrial valleys (Tanaka et al., 1998;
Gulick, 2001), but limited atmospheric precipitation could also
explain the immature development of martian valleys, because
the valley heads are distributed over a wide range of elevations (Mangold and Ansan, 2004; Stepinski et al., 2004). No
detailed analyses are presented here to discriminate between
these modes of formation.
In addition to small valleys, a larger east–west oriented valley of 130-km length and up to 10-km width (DV in Fig. 1)
connects the eastern depression in the Thaumasia highlands
to an impact crater located at the plain–highlands boundary.
MOLA data show that this valley is about 400 m deep on average. This deep valley exhibits two about linear reaches arranged
with an angle of 120◦ (Figs. 1, 2, and 3b). The lower section is
more sinuous with three tributary valleys than the straight upper
section. This linear pattern raises the question of whether the
origin of this valley was volcanic or fluvial. No volcanic flow
is associated with this valley, so it seems unlikely that it was
formed by lava tubes. The deep valley orientation is not concordant with the strike of most normal faults. In contrast, the
valley mouth occurs 250 m below the valley head, indicating
a low longitudinal gradient of 0.15◦ on average. Such a gradient is within the range of fluvial valley gradients (e.g., Leopold
et al., 1992). MOC images cannot provide more information
such as the presence of inner channels because the valley floor
is mantled by transverse aeolian dunes. The deep valley crosscut two N–S normal faults of the plateau. The fact that the deep
valley is not crosscut by any north–south faults shows that it
formed after the main tectonic episodes (Fig. 3b), perhaps in
the Hesperian epoch as proposed by Dohm et al. (2001) (unit
Hch on the geologic map).
Whereas north–south fault have no effects on the shape of
the deep valley, the relatively straight planform of the valley
could be associated with tectonic control along old transverse
east–west faults. The inheritance of faults or fractures is com-
Hydrological system in Thaumasia region, Mars
77
Fig. 1. (a) Location map of the studied region on the south of Tharsis bulge; (b) MOC context mosaic of part of Thaumasia region. G, graben; DV, deep valley; SV,
small valleys.
mon in the development of fluvial valleys (Fig. 4), especially
if the valley extended headward by sapping (Dunne, 1980;
Howard, 1988; Schumm et al., 1995). This process leads to
a headward retreat of valley heads due to groundwater infiltration and undermining of scarps (Pieri, 1980; Dunne, 1980;
Laity and Malin, 1985). The amphitheater head of the deep valley could have its origin in sapping process. The occurrence of
wide but stubby tributaries with theater-shaped heads also supports the sapping process. The geometry of these tributaries is
very similar to those of the terrestrial example (Fig. 4). On the
other hand, surface runoff with overland flow feeding these sapping canyons also exist in this terrestrial example and should
not be neglected in the case of the deep valley. Sapping is usually made possible by precipitation recharge of aquifers and
do not exclude the occurrence of surface runoff (Craddock and
Howard, 2002).
This deep and linear valley displays a geomorphic position
similar to that of terrestrial canyons because it connects the
rift like basin at its source to a less elevated basin at its outlet (Figs. 1–3). Without this deep valley, the elevation of the
divide between the central depression and the western plains
would be about 3700 m, thus more than 900 m higher than
the lower point of the central depression at 2714 m. This topographic setting suggests that the basin to the east of the deep
valley might once have been filled with water and that the valley
was a fluvial connection between this potential source paleolake
and the western impact crater. These hypotheses may be tested
by examining respectively the head and the outlet of the deep
valley.
3. A rift paleolake at the valley head?
Assuming that the current topography is close to the topography at the time of the fluvial activity, the maximum elevation
of the potential lake surface would be about 3700 m which is
the elevation of the divide. A small network of branching val-
Hydrological system in Thaumasia region, Mars
79
Fig. 4. (a) Aerial view of the canyon of Essendilène in SE Algeria. The canyon formed by groundwater sapping and overland flow. Note the presence of a barely
visible valley on the plateau that connects to theater-head valleys which have rectangular shapes typical of control by fractures. This shape is very similar to that of
the central part of the deep valley (in (b)).
morphic features difficult to identify on images because of the
later aeolian mantling. Potential shorelines related to an ocean
in the northern plains have been the subject of intense debates
(e.g., Parker et al., 1993). The southwestern edge of the potential lake is controlled by normal fault scarps bounding grabens
(Fig. 3b), which show interesting similarities with shorelines of
terrestrial rift lakes in Kenya (Fig. 5a). In Landsat image, the
Kenyan fault scarps show locally sinuous contours (arrow A,
Fig. 5a) due to the erosion of the fault scarps. On the edge of
the potential lake, small hills, located adjacent to grabens, suggest that they are the result of erosive processes rather than of
tectonic activity only, because of their irregular jagged shapes
(arrow B, Fig. 5b). They could result from the existence of rectangular bays delimited by normal faults generating grabens
and horsts (arrows B, Fig. 5a). If these landforms are shorelines, then the corresponding elevation of the paleolake would
be around 3400 m.
Paleolake level could also be constrained by the occurrence
of thick sedimentary deposits but no surface property such as
THEMIS nighttime images show specific properties on the potential basin floor. However, Dohm et al. (2001) noted the existence of Hesperian aged smooth terrains of unknown origin
at the southern part of the basin suggesting some resurfacing
processes in this area. Assuming first that the lake occupied the
3400-m level, it would be 150 km long and 70 km wide with a
maximum depth of 650 m and a water volume of the order of
1700 km3 (Fig. 3b). In fact, the depression is divided into two
subbasins: the first in the northwest has a minimum elevation
of 2714 m and the second, to the south, has a minimum elevation of 2989 m. The two depressions separated by a topographic
threshold at about 3250 m which may have isolated the northern depression from the southern one. In a second assumption,
at a lower level at 3200 m, the northern depression would store
about 500 km3 whereas the southern depression connected to
the valley would only contain 60 km3 . At that time, the isolation of the subbasins and level decrease would have reduced the
water supply and stopped the entrenchment of the deep valley.
4. An impact crater paleolake at the valley mouth
The mouth of the deep valley is located on the eastern rim of
a 25-km large impact crater (Fig. 1). THEMIS daytime infrared
images (Figs. 2 and 6a) show two interesting features: (1) the
deep valley ends at a debris fan inside the impact crater and (2)
the western side of the impact crater rim is breached by an outlet
valley which is incised smooth plain and then vanishes to the
west. Using MOLA data (Fig. 6b) we can show that the debris
fan is analogous to terrestrial deltas formed in standing water
and that the outlet valley is due to the overflow of a paleolake
within the impact crater.
80
N. Mangold, V. Ansan / Icarus 180 (2006) 75–87
Fig. 5. (a) Baringo lake in Kenya (Landsat image). (b) Close-up of the southwestern border of the potential Thaumasian rift lake (MOC context image).
4.1. A thick delta at the valley mouth
The debris fan present at the valley mouth exhibits a relatively low-gradient surface bounded by a terminal scarp
(Figs. 6b and 6c). This terminal break in slope takes place 4
km inside the crater inward from the crater rim. The whole debris fan has a difference of elevation of nearly 600 m from its
top surface down to the crater floor, with a maximum elevation
of about 2800 to 2850 m at the top. Fans can exhibit distributary
channels of few tens of meters wide on the top which would indicate long term duration (Malin and Edgett, 2003). However,
the eolian mantle hides small scale features at MOC images
resolution and no erosion have removed this material like in the
Eberswalde crater by Malin and Edgett (2003).
Topographic data can help us to discriminate the nature of
this fan. MOLA data show that the fan’s terminal scarp has a
slope of 11◦ compared to about 2◦ on the fan top surface at
the valley mouth (Fig. 6c). The linearity of the delta profile
(Fig. 6c) measured along a distance of 3 km is uncertain because it could result from interpolation of MOLA points in the
1/128◦ (463 m/cell) DEM. Nevertheless, there are 7 MOLA orbits crossing the delta providing enough points to have a good
approximation of the slope. Moreover, MOC images seem consistent with the general profile with only few effects of subsequent backwasting (Fig. 6d). Indeed, at the base of the terminal
scarp of the fan, a small sharp erosional scarp dissects the fan
(Figs. 6d and 6e). It seems to be less than 50 m high (as seen
on MOLA data and MOC image) and may correspond to backwasting by erosional processes subsequent to deposition. To the
north, there is an horizontal erosional break up inside the crater
wallslope at about 2800 m of elevation (B, Fig. 6d) that could
correspond to a shoreline erosion along the lake border. Additionally, deposits to the south of the fan may be landslides from
the crater rim, but they may also have originated as parts of an
older fan. Nevertheless, it is unlikely that the whole fan would
be a landslide because the upper surface of the fan is at the level
of the valley mouth.
The geometry of the fan is very similar to that of terrestrial
deltas formed under water, i.e., prograding deltas as defined
by Gilbert (1885) where sediments progressively fills the depression (Fig. 7). On Earth, the slope break between the delta
plain (topset) and the delta front (foreset) reflects an abrupt
change from transport by river flow to gravity-driven processes
where the river flow velocity becomes negligible (Leeder, 1999;
Adams et al., 2001). The sharp outer break in slope is especially
consistent with terrestrial Gilbert-type deltas. For example, one
delta in the Lake Mead has a slope that increases from 1◦
to more than 15◦ after the topset–foreset boundary (Fig. 7b).
Lakes in Switzerland have deltas of several hundreds of meters thick that display a foreset slope varying between 14◦ and
27◦ (Adams et al., 2001). In contrast, this geometry is different
from alluvial fans formed subaerially, where the slope break is
concave and occurs at the basin margin. In the martian crater,
the topographic slope break of any potential alluvial fan would
be at the crater border (Fig. 7c). Because the slope break in
the fan is farther away from the crater border, we deduce that
this fan formed under water and corresponds to a paleodelta.
Potential Gilbert-deltas have been identified in imagery at several locations on Mars (e.g., Ori et al., 2000a) but few examples
have been studied with detailed topography. Here the paleodelta
Hydrological system in Thaumasia region, Mars
81
Fig. 6. (a) Mosaic of two THEMIS IR images of the crater at the mouth of the valley. The arrow D shows the delta at the deep valley mouth, C shows the outlet
channel that flooded the plains and B shows the border of the crater eroded at the same level than the delta plain. (b) MOLA DEM of the same area. The black line
corresponds to the cross-section of figure (c). (d) High resolution MOC mosaic and structural schema (e) with MOLA elevation contour lines superimposed.
would represent a minimum of 15 km3 of material deposited by
the valley inflow.
The curvature of terrestrial delta slopes can be linear, sigmoidal or exponential depending on the processes involved
82
N. Mangold, V. Ansan / Icarus 180 (2006) 75–87
Fig. 7. (a) Theoretical cross-section of a delta fan (adapted from Birkeland and
Larson, 1989). (b) Terrestrial example of a delta in Las Vegas Bay, Lake Mead,
Nevada (data from Twichell et al., 2001). (c) Theoretical cross-section of an
alluvial fan.
(e.g., Adams et al., 2001). Exponential and sigmoidal slopes
are the result of gravity instabilities and erosional modifications typical in submarine environment where specific effects
like storms, tides and level changes produce large instabilities such as turbidites (Schlager and Adams, 2001). In contrast, linear profiles represent deposits lying at the angle of
repose in absence of strong instabilities such as in quite lake.
The 11◦ linear frontal slope of the fan inside the impact crater
(Fig. 6c) could reflect the angle of repose of the material. The
low value of 11◦ is a strong argument in favor of deposition
under water because the angles of repose of all dry materials
range from 25◦ to 35◦ because of friction properties (Selby,
1982). Angles of repose under water depend mainly on the
size of the particles and the high pore pressures reduces the
stable talus angle by half (Selby, 1982). Terrestrial examples
show that materials dominated by small particles like silt and
clay (less than 62 µm) are dominant in deltas at 14◦ of slope
(Adams et al., 2001). Similar fine-grained clastic grains could
thus compose the material of the delta in the impact crater
(Fig. 6).
4.2. An overflow channel west of the crater paleolake
The outlet valley to the west of the crater paleolake has a
very different shape than the deep valley or any of the smaller
valleys in the area. This valley is in fact a single channel of
3 km wide near its head but it progressively widens and disappears on the flat plains only 20 km to the west (Figs. 2, 3).
Such shape is characteristic of a quick release of water in the
same way as large-scale outflow channels (e.g., Baker, 1982),
especially because of the lack of tributary and the strong widening during downflow. Thus, this channel is likely formed by a
brief flood rather than through prolonged erosion forming valley networks. This discharge could be due to the overflow of
the lake into the surrounding plains. The topography can be
used to test this hypothesis. The exit breach is located on the
rim of the crater at an elevation of about 2750 m (Fig. 6b).
This part of the crater rim is entrenched relative to other sections of the fresh crater rim, which stand at about 3000 m of
elevation. At the mouth of the deep valley, east of the crater,
the maximum elevation of the delta topset plain is of 2800
to 2850 m (uncertainty due to MOLA spatial resolution). Assuming that the crater was once filled by water to the delta
surface, this uppermost water level was thus 50 to 100 m
more elevated than the broken rim at the outlet channel at the
exit breach. On the other hand, if the lake was filled up to
the 2800 m level (higher than the surrounding plains at 2600
m of elevation) ponded water likely infiltrated into the crater
rim. The saturation of groundwater paths may then have produced the local collapse of the rim inducing a flood out of the
crater.
If the outlet channel was created by the crater lake overflow, then the volume of water must have been larger than this
eroded volume. The outlet channel is 20 km long and 50 m
deep on average and produced an eroded volume of 4.3 km3
which we can take as a minimum for the water amount. On
the other hand, the crater lake would have released a layer of
water of 50–100 m thick, thus a total flood volume of 25–
50 km3 . This value is in agreement with the minimum eroded
volume of 4.3 km3 if the proportion of water in the flow would
have been of 10–20% by volume. Such value is in agreement
with concentrated flows of rock–water mixture (Corominas et
al., 1996). At the western edge of the outlet channel, an elliptical area resembling to a depositional fan is visible on the
nighttime THEMIS image (Fig. 2b) but this possible deposition zone is too small and thin to be evaluated quantitatively.
The paleolake level deduced from the delta plain elevation
and the outlet breach level shows that the crater was filled by
more than 500 m of water at the time of the overflow, implying
a total volume of water of 250 km3 . This value is a minimum
if the crater was once deeper and that sediments progressively
filled a part of the interior of the crater. Indeed, fresh craters
of 25 km in diameter have depths reaching 1.5–2 km according to empirical laws (Garvin et al., 2002). With 700 m deep,
this crater is thus not fresh and may be partially filled by sediments as many ancient craters (Craddock and Howard, 2002).
In this case, more than 800 m of sediments corresponding to a
maximum volume of sediments of 400 km3 could remain in the
crater. On the other hand, the total volume of material eroded
along the 120 km of the deep valley is in the range of 150 km3 .
By comparison, the volume of the delta is ten times less, suggesting that most of the material eroded is present beneath the
current crater floor.
Hydrological system in Thaumasia region, Mars
5. Discussion
5.1. Duration of the hydrological activity and origin of the
water
Crater lakes may have formed during brief episodes if they
are due to rapid groundwater discharge. Indeed, outflow channels induced by volcano-tectonic, geothermal or subsurface water activity such as Mangala Valles or Athabasca Vallis (Head
et al., 2003; Burr et al., 2002) can produce floods that fill
crater floors with sediments. The deep valley and the crater
lake discussed here were once interpreted as “erosional remnants formed by disruption of terrain by groundwater release
and tectonic activity or mass wasting” (Dohm et al., 2001). This
would suggest either an origin similar to Mangala Valles, but
with a smaller size, or similar to Maádim Vallis which formed
from the discharge of lakes in the source area (Irwin et al.,
2002). However, the presence of the Gilbert-delta at the valley outlet is a strong argument against a catastrophic release
of liquid water for the deep valley formation. Outflow channels have no Gilbert-delta at their mouth because they do not
involve a continuous filling of sediments by water and have
a strong velocity at valley mouth. The presence of the delta
implies the progressive accumulation of sediments, thus fluid
flows over geologically significant period of time. Moreover,
the deep valley does not display any braided channels, fluvial
dunes or tear-dropped islands which usually characterize the
quick and high liquid flux of outflow channels. It is nevertheless
possible that an outflow channel incised part of the deep valley
at the beginning of the connection between the upper and the
lower lake and that a gentle surface runoff continued the erosion processes and formed the delta-like fan. On the contrary,
the outlet valley to the west of the crater correspond to a unique
episode of quick water release without fluvial style of runoff.
Nevertheless, a progressive filling of the crater lake does not
mean that the fluvial activity occurred over geologically long
periods of time. Terrestrial deltas may form relatively quickly
(e.g., Adams et al., 2001). In the absence of strong tides of
wave effects, some terrestrial marine deltas exhibit deposition
rates of several centimeters per year especially if they are fed
by large sediment-laden rivers such as the Ganges or the Mississippi (Mangold and Ansan, 2004). Lacustrine deltas exhibit
thick deposition of clastic sediments that can give information
on their deposition rates. Since the last ice ages, several hundred
meters of sediment have been deposited in most lakes (Adams
et al., 2001). Artificially damned rivers permits estimates of the
sedimentation of deltas up to rates of, for example, deltaic accumulations of 40 m in 80 years occurred at Lake Constance
along the Rhine (Müller, 1966), and 20 m of sediments accumulated in the delta of the Las Vegas Bay in Lake Mead over
60 years (Twichell et al., 2001). The total volume deposited can
be used to estimate durations given some assumptions on discharge rate. The minimum volume of the martian delta can be
estimated at 15 km3 . By comparison, the Colorado river delta
at Lake Mead formed in 60 years is only 0.0019 km3 to date.
Lake Brienz, Switzerland, has a similar volume to the martian
lake with 5.2 km3 of sediments deposited during the Quaternary
83
epoch (Adams et al., 2001). This shows that several thousands
of years are required to deposited such amount of material under conditions corresponding to interglacial climate in the midlatitudes on Earth. This estimate of duration can be taken as a
strict minimum for the formation of such a thick delta on Mars
because of the much more wet climate on Earth. Even if we assume that Mars was more wet in the early periods of time, it
would unlikely reach precipitations of Switzerland of typically
1–2 m/yr in mountains (Adams et al., 2001). Inversely, maximum durations for the delta formation are difficult to estimate
because the delta may have formed by slow accumulation of
transient flows under an arid climate. Nevertheless, in this case,
the lake should have remained at the level of the delta plain
during long time which is unlikely to occur over durations of
million of years because evaporation would predominate over
water influx.
Lakes levels on Earth are controlled by the water table (e.g.,
Leeder, 1999), especially if the lakes exist over a significant
duration. The fact that the deep valley emerges inside a crater
that includes the lowest elevation in the region (i.e., 2200 m
compared to 2600 m in the plains) is certainly not a coincidence. Groundwater may have partially filled this crater even
before the deep valley connected the upper lake to the lower
lake. Groundwater was also likely responsible of the breaching of the crater rim which permitted the flooding of the nearby
plain. For the potential rift lake in the eastern part of the studied
region, the presence of groundwater was necessary to sustain
the presence of a perennial lake. The fact that this depression is
surrounded by mountains of 5–6 km elevation is consistent with
this possibility. It is nevertheless difficult to know if the deep
valley was mainly supplied by groundwater as some characteristics of its morphology suggest may indicate, or if the potential
upper lake was a major contributor. It is also possible that the
source of the deep valley resulted only from surface and subsurface flows collected from the interior rift basin without forming
a lake.
The presence of a large lake or even a spring discharge raises
questions regarding the origin of the water. Small branching
valleys are identified in the region, but only few are located
around the head of the deep valley or the lake boundary. Data
also suggest that subsurface flows were probably important.
The duration and volume of water involved with this hydrologic system are large and difficult to be explained by local
hydrothermal systems or by seepage due to tectonic activity
alone. Melting from glaciers could provide liquid water (Carr
and Head, 2003), but no glacial valley, flutes or grooves indicating any glacier movement, or any potential moraines are
observed in this area. Rainfall, even in an arid climate, could
explain high water volumes. Nevertheless, if fluvial erosion is
triggered mainly by rainfall, the fault scarps should also be
eroded by runoff. This seems not to be the case for most of
the scarps, even the oldest ones. However, it should be noted
that faults of the African rift (Fig. 5) are devoid of apparent
runoff gullies in the left part of the image. This is especially
the case in a dry climate when permeable lava flows result in
high infiltration rates relative to runoff. Sapping is usually also
made possible by precipitation recharge of aquifers (Craddock
84
N. Mangold, V. Ansan / Icarus 180 (2006) 75–87
and Howard, 2002). This issue of the water source remains nevertheless largely unresolved, as it is for many ancient martian
landforms.
5.2. Scenario of lakes formation
Our study shows that the interpretation of a crater lake with
the delta and the overflow breach is well constrained by imaging and topographic data. The deep valley probably requires
flows of relatively long duration for its formation, development
of tributaries and formation of the delta at its mouth, although
that the relative contributions of surface and subsurface flows
remains uncertain. The presence of the upper lakes is consistent
with the necessity of a water source to erode the deep valley,
given that this large topographic depression is only connected
to lower plains by this deep valley. Shorelines may indicate past
levels of this lake but they are ambiguous landforms, and the
evidence for this upper paleolake is less conclusive than for a
lake in the lower breached crater. Using these interpretations we
can propose a scenario for the development of these landforms
(Fig. 8).
(1) The central depression was closed and possibly filled by
water up to the level of 3400 m with possible shorelines.
The water table was high in this basin relative to the western plains which resulted in groundwater flow to lower areas. The impact crater at that time might have been filled
with water at that time, although no evidence of flooding during that epoch remains. At that time, the deep valley began to extend headward by groundwater sapping.
This mechanism why the valley debouched into the crater
and not to the surrounding plains. The crater floor was
deeper than the surrounding which made it a natural site
for groundwater discharge.
(2) After the fluvial connection was established from the upper lake to the lower lake in the crater, more sediments
were provided to fill the crater. A quick release from the
upper lake is possible at that time but no landforms from
such episodes are preserved. The upper lake level may have
beginning to decline because of the greater discharge that
would be associated with a surface outlet. The delta begins
to built mainly during that episodes of great fluvial connectivity.
(3) In a next step, the upper lake was likely less elevated than
it was previously with a level near the elevation of the deep
valley head. The deep valley was at its maximum of incision and the delta had formed and was prograding inside
the crater. At the level of 2800 m, the lower lake water level
was higher than the surrounding plains inducing a breach in
the crater rim and an overflow in the plains.
fluvial activity as Noachian in age (Malin and Edgett, 2000).
Nevertheless, crater lakes in the Ismenius Lacus and Memnonia region have been recognized on the basis of potential
deltas of Hesperian age (Ori et al., 2000a). Hesperian or Early
Amazonian ages were proposed for some crater lakes in the
highlands based on small crater populations on the crater lake
floor (Cabrol and Grin, 1999, 2001). Ages are obtained for only
30 of 179 craters from crater counts at Viking resolution. This
spatial scale limits the interpretation of age because of the low
craters statistics due to the small area sampled. This is also the
case for the crater lake of this study. Nevertheless, the hydrological activity identified here postdates the Noachian terrain in
which the valleys are incised. This system has been dated to the
end of the Early Hesperian by Dohm et al. (2001) and this age
is likely a minimum. First, the deep valley is younger than all
tectonic faults which are dated up to the Hesperian period because some of them cut the Hesperian plains. Second, the outlet
valley, west of the crater lake, is superimposed on the smooth
plains dated at the Noachian–Hesperian boundary. Our study
therefore shows that the hydrological activity occurred after the
Noachian–Hesperian boundary. No upper stratigraphic bound
can be given due to the lack of younger structures.
The Hesperian epoch is usually interpreted as cold and dry
with a thick permafrost layer (e.g., Carr, 1996). Lakes can be
blanketed by ice and be stable at cold temperatures during
long periods of time. In fact, the subsurface flows may have
filled the crater lake by subsurface processes only at the beginning of the process, possibly in a permanently frozen climate, if seepage from geothermal sources was possible. However, the presence of the deep valley, with different tributaries
pleads in favor of surface flows. Indeed, sapping is a subsurface process until the water comes at the surface at the
head of the valley. After the valley head, the flow is subaerial,
which correspond here to several tens of kilometers of flows
from the different heads of the deep valley and its tributaries.
Thus, it is difficult to explain how the valleys can feed the
crater lake up to 600 m depth in a permanently frozen climate. The whole suite of flows from the deep valley formation
to the overflow into surrounding plains is therefore easier to
explain if liquid water was stable at the surface, at least seasonally in a cold climate, or episodically in an arid one. Surface
flows and potential lakes of Late Hesperian epoch are identified in region like Valles Marineris (Mangold et al., 2004;
Quantin et al., 2004) raising a question as to the nature of the
Hesperian climate. An arid climate with episodic atmospheric
precipitation could have made a progressive transition between
a relatively wet Noachian period and a dry Amazonian period,
which may explain Hesperian lakes (Craddock and Howard,
2002). An alternative solution is a transient episode of fluvial
activity in valleys following of the outflow channels activity
(Baker et al., 1991).
5.3. Age of the hydrological activity
6. Conclusion
The possibility of standing bodies of liquid water on Mars
has been demonstrated by the observation of a delta with meandering channels in inverted relief (Malin and Edgett, 2003).
Most investigators interpret potential paleolakes and associated
Using visible and thermal IR imaging at different resolutions, along with topographic data, we detail an occurrence of
hydrologic activity in the southwestern part of the Thaumasia
Hydrological system in Thaumasia region, Mars
85
Fig. 8. Scenario of evolution of the lake system. On top, synthetic map with the key features. Below are three stages of evolution. See text for explanations.
highlands, south of the Tharsis bulge. A deep valley was incised into the highlands and delivered water to a paleolake in
an impact crater as shown by the similarity of a fan on the eastern side of the crater to terrestrial Gilbert-type deltas. An exit
breach on the rim of the crater attests to the overflow of the
crater paleolake. A larger depression at the contributing valley
head might have provided surface water to the crater paleolake.
Stratigraphic relationships indicate an age in the Hesperian period, later than the usual Noachian–Hesperian boundary that is
usually suggested as the end of major fluvial activity on Mars.
Surface flows nevertheless might have had relatively short duration, as such lakes can form in periods of several thousand
to tens of thousand of years in climates corresponding to midlatitudes regions on Earth. The occurrence of late warm episode
of surface flows and water detention in lakes is important for the
preservation of depositional forms whereas lakes in general are
interesting for any exobiological study on Mars (e.g., Ori et al.,
2000b).
Acknowledgments
The authors greatly acknowledge D. Burr and R. Irwin for
their detailed and very helpful reviews and the NASA, USGS,
MSSS and ASU web facilities for the use of MOC, MOLA and
THEMIS data. This work is supported by the PNP (Prog. Nat.
de Planétologie) of INSU (Institut National Sciences de l’Univers) of France and the MAGE (Mars Geophysical NEtwork)
network of the European Community Improving Human Potential Program (contract RTN2-2001-00414).
References
Adams, E.W., Schlager, W., Anselmetti, F., 2001. Morphology and curvature of
delta slopes in Swiss lakes: Lessons for the interpretation of clinoforms in
seismic data. Sedimentology 48, 661–679.
Ansan, V., Mangold, N., 2003. Warrego Valles revisited using MGS and
Odyssey data: Valleys formed by precipitations? In: VIth Int. Conf. on
Mars, Pasadena, USA.
Ansan, V., Mangold, N., 2005. New map and morphometry of Warrego Valles
from MGS and Mars Odyssey data. Planet. Space Sci. Submitted for publication.
Baker, V., 1982. The Channels of Mars. Univ. of Texas Press, Austin.
Baker, V.R., Strom, R.G., Gulick, V.C., Kargel, J.S., Komatsu, G., Kale, V.S.,
1991. Ancient oceans, ice sheets and the hydrological cycle on Mars. Nature 352, 589–594.
Birkeland, P., Larson, E.E., 1989. Putnam’s Geology, fifth ed. Oxford Univ.
Press, New York. 648 p.
86
N. Mangold, V. Ansan / Icarus 180 (2006) 75–87
Burr, D.M., Grier, J.A., McEwen, A.S., Keszthelyi, A.P., 2002. Repeated aqueous flooding from the Cerberus Fossae: Evidence for very recently extant
deep groundwater on Mars. Icarus 159, 53–73.
Cabrol, N.A., Grin, E.A., 1999. Distribution, classification and ages of martian
impact crater lakes. Icarus 142, 160–172.
Cabrol, N.A., Grin, E.A., 2001. The evolution of lacustrine environments on
Mars: Is Mars only hydrologically dormant? Icarus 149, 291–328.
Cabrol, N.A., Grin, E.A., 2002. Overview on the formation of paleolakes and
ponds on Mars. Global Planet. Change 35, 199–219.
Carr, M.H. (Ed.), 1996. Water on Mars. Oxford Univ. Press, New York. 229 p.
Carr, M.H., Head III, J.W., 2003. Basal melting of snow on early Mars: A
possible origin of some valley networks. Geophys. Res. Lett. 30 (24),
doi:10.1029/2003GL018575. 2245.
Christensen, P.R., and 22 colleagues, 2003. Morphology and composition of
the surface of Mars: Mars Odyssey THEMIS results. Science 300 (5628),
2056–2061.
Clifford, S.M., 1993. A model for the hydrologic and climatic behavior of water
on Mars. J. Geophys. Res. 98, 10973–11016.
Colaprete, A., Toon, O.B., 2000. The radiative effects of martian water ice
clouds on the local atmospheric temperature profile. Icarus 145, 524–532.
Corominas, J., Remondo, J., Farias, P., Estevao, M., Zezere, J., Dias de Teran,
J., Dikau, R., Schrott, L., Moya, J., Gonzalez, A., 1996. Debris flow. In:
Dikau, R., Brundsen, D., Schrott, L. (Eds.), Landslide Recognition. Wiley,
New York, pp. 161–180.
Craddock, R.A., Howard, A.D., 2002. The case for rainfall on a warm, wet early
Mars. J. Geophys. Res. 107 (E11). 5111.
Craddock, R.A., Maxwell, T.A., 1990. Resurfacing of the martian highlands
in the Amenthes and Tyrrhena region. J. Geophys. Res. 95 (B9), 14265–
14278.
De Hon, R.A., 1992. Martian lake basins and lacustrine plains. Earth Moon
Planets 56, 95–122.
Dohm, J.M., Tanaka, K.L., 1999. Geology of the Thaumasia region, Mars.
Planet. Space Sci. 47, 411–431.
Dohm, J.M., Tanaka, K.L., Hare, T.M., 2001. Geologic map of the Thaumasia
region of Mars. U.S. Geological Survey Map I-2650.
Dunne, T., 1980. Formation and controls of channel networks. Prog. Phys.
Geog. 4, 211–240.
Forget, F., Pierrehumbert, R.T., 1997. Warming early Mars with carbon dioxide
clouds that scattered infrared radiation. Science 278, 1273–1276.
Garvin, J.B., Sakimoto, S.E.H., Frawley, J.J., Schnetzler, C., 2002. Global geometric properties of martian impact craters. Lunar Planet. Sci. 33. Abstract
1255.
Gilbert, G.K., 1885. The topographic features of lake shores. U.S. Geol. Surv.
Annu. Rep. 5, 75–123.
Grant, J.A., 2000. Valley formation in Margaritifer Sinus, Mars, by
precipitation-recharged ground-water sapping. Geology 28, 223–226.
Gulick, V.C., 2001. Origin of the valley networks on Mars: A hydrological perspective. Geomorphology 37 (3–4), 241–268.
Head, J.W., Wilson, L., Mitchell, K.L., 2003. Generation of recent massive
water floods at Cerberus Fossae by dike emplacement, cryospheric cracking, and confined aquifer groundwater release. Geophys. Res. Lett. 30 (11).
1577.
Howard, A.D., 1988. Saping features of the Colorado plateau. In: Howard,
A.D., Kochel, R.C., Holt, H.E. (Eds.), A Comparative Planetary Geology
Field Guide. NASA SP-491, p. 108.
Irwin, R.P., Maxwell, T.A., Howard, A.D., Craddock, R.A., Leverington, D.W.,
2002. A large paleolake basin at the head of Maádim Vallis, Mars. Science 296 (5576), 2209–2212.
Kasting, J.F., 1991. CO2 condensation and the climate of early Mars. Icarus 94
(1), 1–13.
Laity, J.E., Malin, M.C., 1985. Sapping processes and the development of
theater-headed valley networks in the Colorado plateau. Geol. Soc. Am.
Bull. 96, 203–217.
Leeder, M., 1999. Sedimentology and Sedimentary Basins. Blackwell Sci., Oxford. 592 p.
Leopold, L.B., Wolman, M.G., Miller, J.P., 1992. Fluvial Processes in Geomorphology. Dover, New York. 524 p.
Malin, M.C., Edgett, K.S., 2000. Sedimentary rocks of early Mars. Science 290,
1927–1937.
Malin, M.C., Edgett, K.S., 2003. Evidence for persistent flow and aqueous sedimentation on early Mars. Science 302, 1931–1934.
Malin, M.C., Carr, M.H, Danielson, G.E., Davies, M.E., Hartmann, W.K., Ingersoll, A.P., James, P.B., Masursky, H., McEwen, A.S., Soderblom, L.A.,
Thomas, P., Veverka, J., Caplinger, M.A., Ravine, M.A., Soulanille, T.A.,
Warren, J.L., 1998. Early views of the martian surface from the Mars Orbiter Camera of Mars Global Surveyor. Science 279, 1681–1685.
Mangold, N., Ansan, V., 2004. Long term stability of liquid water on early
Mars from the identification of lakes and deltas. In: 1st Eur. Geophys. Union
Symposium, Nice. Abstract.
Mangold, N., Quantin, C., Ansan, V., Delacourt, C., Allemand, P., 2004.
Evidence for precipitation on Mars from dendritic valleys in the Valles
Marineris area. Science 305 (5680), 78–81.
Michels, K.H., Kudrass, H.R., Hübscher, C., Suckow A., Wiedlicke, W., 1998.
The submarine delta of the Ganges–Brahmaputra: Cyclone-dominated sedimentation patterns. Mar. Geol. 149, 133–154.
Mischna, M.A., Kasting, J.F., Pavlov, A., Freedman, R., 2000. Influence of carbon dioxide on early martian climate. Icarus 145, 546–554.
Müller, G., 1966. The new Rhine delta in Lake Constance. In: Shirley, M.L.,
Ragsdale, J.E. (Eds.), Deltas in Their Geologic Framework. Houston Geologic Society, Texas, pp. 107–124.
Needham, H.D., Choukroune, P., Cheminée, J.L., LePichon, X., Francheteau,
J., Tapponnier, P., 1976. The accreting plate boundary Ardoukoba rift (NE
Africa) and the oceanic rift valley. Earth Planet. Sci. Lett. 28 (3), 439–453.
Newson, H.E., Brittelle, G.E., Hibbits, C.A., 1996. Impact crater lakes on Mars.
J. Geophys. Res. 101 (E6), 14951–14955.
Ori, G.G., Marinangelli, L., Baliva, A., 2000a. Terraces and Gilbert-type
deltas in crater lakes in Ismenius Lacus and Memnonia, Mars. J. Geophys.
Res. 105 (E7), 17629–17641.
Ori, G.G., Marinangelli, L., Komatsu, G., 2000b. Martian paleolacustrine environments and their geological constraints on drilling operations for exobiological research. Planet. Space Sci. 48, 1027–1034.
Parker, T.J., Gorsline, D.S., Saunders, R.S., Pieri, D.C., Schneeberger, D.M.,
1993. Coastal geomorphology of the martian northern plains. J. Geophys.
Res. Planets 98 (E6), 11061–11078.
Pieri, D.C., 1980. Martian valleys: Morphology, distribution, age and origin.
Science 210, 895–897.
Pollack, J.B., Kasting, J.F., Richardson, S.M., Poliakoff, K., 1987. The case for
a wet and warm climate on early Mars. Icarus 71, 203–224.
Quantin, C., Allemand, P., Mangold, N., Dromart, G., Delacourt, C., 2004. Paleolake and Dense Valley networks originated by precipitations within Melas
Chasma, Valles Marineris. In: 3rd Early Mars Conference, Jackson lake.
Abstract 8010.
Sagan, C., Toon, O.B., Gierasch, P.J., 1973. Climate changes on Mars. Science 181, 1045–1049.
Schlager, W., Adams, E.W., 2001. Model for the sigmoidal curvature of submarine slopes. Geology 29 (10), 883–886.
Schultz, R.A., Tanaka, K.L., 1994. Lithospheric-scale buckling and thrust structures on Mars: The Coprates rise and south Tharsis ridge belt. J. Geophys.
Res. 99 (E4), 8371–8385.
Schumm, S.A., Mosley, M.P., Weaver, W.E., 1987. Experimental Fluvial Geomorphology. Wiley, New York. 413 p.
Schumm, S.A., Boyd, K.F., Wolff, C.G., Spitz, W.J., 1995. A ground-water
sapping landscape in the Florida Panhandle. Geomorphology 12, 281–
297.
Selby, M.J., 1982. Hillslope Materials and Processes. Oxford Univ. Press, Oxford. 466 p.
Smith, D.E., Zuber, M.T., Solomon, S.C., Phillips, R.J., Head, J.W., Garvin,
J.B., Banerdt, W.B., Muhleman, D.O., Pettengill, G.H., Neumann, G.A.,
Lemoine, F.G., Abshire, J.B., Aharonson, O., Brown, C.D., Hauck, S.A.,
Ivanov, A.B., McGovern, P.J., Zwally, H.J., Duxbury, T.C., 1999. The
global topography of Mars and implications for surface evolution. Science 284, 1495–1503.
Stepinski, T.F., Collier, M.L., McGovern, P.J., Clifford, S.M., 2004. Martian
geomorphology from fractal analysis of drainage networks. J. Geophys.
Res. 109 (E2). E02005.
Hydrological system in Thaumasia region, Mars
Tanaka, K.L., Dohm, J.M., Lias, J.H., Hare, T.M., 1998. Erosional valleys in the
Thaumasia region of Mars: Hydrothermal and seismic origins. J. Geophys.
Res. 103 (E13), 31407–31420.
87
Twichell, D.C., Cross, V.A., Rudin, M.J., Parolski, K.F., Rendings, R.R., 2001.
Surficial geology and distribution of post-impoundment sediment in Las
Vegas Bay, Lake Mead. USGS, Open Report 01-070.