PUBLICATIONS
Journal of Geophysical Research: Planets
RESEARCH ARTICLE
10.1002/2015JE004936
Key Points:
• The formation of the South Tharsis
Ridge Belt may provide information
on the early evolution of Tharsis and
early Mars
• The shape, size, and separation of the
ridges resemble the extensional Basin
and Range Province on Earth
• Both the morphology and associated
crustal thinning support an extensional
origin for the ridges
Supporting Information:
• Supporting Information S1
Correspondence to:
E. Karasözen,
[email protected]
Citation:
Karasözen, E., J. C. Andrews-Hanna,
J. M. Dohm, and R. C. Anderson (2016),
The formation of the South Tharsis Ridge
Belt: Basin and Range-style extension
on early Mars?, J. Geophys. Res. Planets,
121, 916–943, doi:10.1002/
2015JE004936.
Received 6 SEP 2015
Accepted 6 MAY 2016
Accepted article online 13 MAY 2016
Published online 4 JUN 2016
©2016. American Geophysical Union.
All Rights Reserved.
KARASOZEN ET AL.
The formation of the South Tharsis Ridge Belt: Basin
and Range-style extension on early Mars?
Ezgi Karasözen1, Jeffrey C. Andrews-Hanna1,2, J. M. Dohm3, and R. C. Anderson4
1
Department of Geophysics and Center for Space Resources, Colorado School of Mines, Golden, Colorado, USA, 2Southwest
Research Institute, Boulder, Colorado, USA, 3The University Museum, The University of Tokyo, Tokyo, Japan, 4Jet Propulsion
Laboratory, California Institute of Technology, Pasadena, California, USA
Abstract The South Tharsis Ridge Belt (STRB) is located along the northeastern edge of Terra Sirenum and
partially surrounds the southwestern part of Tharsis in an arc. It consists of 29 large ridges separated by
distances 130 to 260 km, with average relief of 1.5 km above the surrounding plains. Because the STRB is
among the oldest tectonic features associated both spatially and developmentally with Tharsis, it may
provide key information on the early evolution of Tharsis and possibly pre-Tharsis processes. Earlier studies
concluded that the ridges formed through compressional tectonism. However, the shape, size, and
separation of the ridges support the interpretation that the STRB resembles the extensional Basin and Range
Province on Earth. Both regions are characterized by series of parallel mountain ranges separated by broad
valleys. In this study, we evaluate both extensional and compressional hypotheses for the origin of the ridges
using evidence from topography, deformed craters, crustal thickness models, and strain modeling. Though
neither interpretation explains all aspects of the ridges, the topography of the ridges and crustal thinning
associated with the western part of the ridge belt support an extensional origin. Strain models predict that
Basin and Range-style wide rifting would be expected for early Martian conditions. This extension may have
been initiated by plume-induced uplift in the early stages of Tharsis formation, but the large amount of
extensional strain inferred in the western STRB must have been accommodated by compression elsewhere,
possibly in the heavily deformed craters of western Terra Sirenum.
1. Introduction
The tectonic histories of the planets are coupled to their geodynamic evolution. On Earth, most of the tectonic
features are associated with plate tectonics and a mobile lithosphere. However, hypothesized plate tectonism
on Mars, including reported geological and geophysical supporting evidence [Sleep, 1994; Dohm et al., 2001a,
2002, 2013, 2015; Anguita et al., 2001; Fairén et al., 2002; Fairén and Dohm, 2004; Connerney et al., 2005; Baker
et al., 2007; Yin, 2012a, 2012b], remains controversial. For example, geophysical evidence suggests that the crustal
dichotomy likely formed in a giant impact [Andrews-Hanna et al., 2008] rather than through plate tectonic
processes [e.g., Sleep, 1994]. Most of the tectonic features from the Late Noachian onward are consistent with
the interaction of global contraction and membrane-flexural deformation of the lithosphere in response to loading by Tharsis and other volcanic constructs [Banerdt et al., 1992; Banerdt and Golombek, 2000; Anderson et al.,
2001; Andrews-Hanna et al., 2008]. Among these Late Noachian and later tectonic features are the pervasive
wrinkle ridges and graben swarms, which are characterized by small strains and faulting in an otherwise laterally
continuous lithosphere (i.e., not requiring plate tectonics). However, a subset of ancient tectonic features of
pre-Noachian to Early Noachian age cannot be easily explained by these processes, including rift zones
[Grott et al., 2007; Hauber et al., 2010], Valles Marineris [Andrews-Hanna, 2012a, 2012b, 2012c; Yin, 2012a],
ancient structurally controlled basins in Terra Cimmeria [Fairén et al., 2002], and mountain ranges including
the Thaumasia highlands [Dohm and Tanaka, 1999; Dohm et al., 2001b, 2015; Fairén and Dohm, 2004;
Montgomery et al., 2009; Nahm and Schultz, 2010]. These ancient features display significant localized
strains that are difficult to reconcile with the Martian lithosphere behaving as a perfectly stagnant lid.
Investigating the possible mechanism(s) responsible for the formation of these features through geologic
mapping and geophysical assessment may help shed light on the nature of early Martian geodynamics.
One of these ancient tectonic features is the South Tharsis Ridge Belt (STRB) (Figure 1), mapped and interpreted by Schultz and Tanaka [1994] to consist of a system of 29 ridges forming an arcuate shape around
southern Tharsis. The ridges are separated by distances of ~150–500 km and have reliefs of ~1.5 km above
the surrounding plains. The STRB is among the oldest tectonic features associated with Tharsis, but the
FORMATION OF THE SOUTH THARSIS RIDGE BELT
916
Journal of Geophysical Research: Planets
10.1002/2015JE004936
Figure 1. MOLA topography map [Smith et al., 2001] in a cylindrical projection, showing the western part of the South
Tharsis Ridge Belt in northern Terra Sirenum (white box; Figure 2a) and western Terra Sirenum (black box; Figure 8). Also
shown is the location of Amenthes Rupes (yellow box; Figure 2d). Note the nonlinear elevation scale and the saturation of
the topography at 4.69 km, chosen so as to emphasize the structures of interest. Image made with 128 pixels/degree
gridded MOLA data.
mechanism of its formation is unclear. Based on photogeologic interpretations, Schultz and Tanaka [1994]
suggested that the STRB formed by a combination of a buckling instability and thrust faulting related to early
Tharsis development. However, we note here that the geomorphology and morphometry of the ridges of the
western part of the south Tharsis ridge belt in northern Terra Sirenum resemble the parallel mountain ranges
of the Basin and Range Province of the western United States [Anderson et al., 2012]. Both provinces are made
up of sets of quasi-parallel steep-sided ranges separated by broad flat basins. The ranges of the Basin and
Range of the western United States are bounded by normal faults that together with ductile deformation
of the lower crust accommodate regional extensional strain of approximately 100% [Hamilton and Myers,
1966]. Similarities in the shape, size, and separation of the basins and ranges of the western STRB and the
Basin and Range, as detailed in this study, suggest a very different interpretation for these ridges from the
previous compressional model, with significant implications for the early evolution of Mars.
We here examine the geophysical evidence for the nature and origin of the western STRB, particularly on
those ridges located in the northern part of the Terra Sirenum region. We focus on the western STRB since
the ridges farther south and east extend higher up on the Tharsis rise and appear to be more heavily modified
by Tharsis volcanism. In related work, Anderson et al. [2012] investigate this region through geologic mapping. In this study, we evaluate both hypothesized extensional and compressional origins for the formation
of the ridges through topographic profiles, deformed impact craters, crustal thickness models, and strain
modeling. We compare the western STRB (Figures 2a and 2b) with the Basin and Range Province of Earth
(Figure 2c) as an extensional analog and the Amenthes Rupes thrust fault on Mars (Figure 2d) as a compressional analog. In section 2, we discuss these analogs and their tectonic environments. In section 3, we compare topographic profiles of the western STRB to the Basin and Range Province and to Amenthes Rupes and
use these profiles to constrain the western STRB fault geometry. In section 4, the strain across the western
STRB is evaluated using impact craters and crustal thickness models and then possible strain mechanisms
are discussed. In section 5, strain modeling is carried out to understand the implications that Basin and
Range-style rifting would have for conditions on Noachian Mars. In section 6, possible scenarios for the
tectonic and geodynamic evolution of the STRB are discussed. Though no one model explains all aspects
of the ridges, the topography of the ridges and crustal thinning associated with the ridge belt support an
extensional origin, which has significant implications for the incipient development of Tharsis and possibly
pre-Tharsis phases of Mars’ evolution.
2. Compressional and Extensional Analogs to the South Tharsis Ridge Belt
2.1. Compressional Analog: Amenthes Rupes
We first discuss possible compressional analogs to the STRB. During horizontal compression or extension, the
lithosphere can become unstable, resulting in buckling or necking. The STRB was interpreted as a set of
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
917
Journal of Geophysical Research: Planets
10.1002/2015JE004936
Figure 2. Topographic maps of (a, b) the South Tharsis Ridge belt in the western Terra-Sirenum region of Mars, (c) the Basin
and Range Province on Earth, and (d) Amenthes Rupes on Mars. Figure 2a shows the location (white box) of a close-up of a
prominent north-trending ridge located near the source region of Mangala Valles detailed in Figure 2b, as well as the
transect lines (A-A′, B-B′, and C-C′) corresponding to the topographic profiles shown in Figure 3 and (1-1′, 2-2′) to the profiles
shown in Figure 4. The locations of investigated impact craters (letters with boxes) are shown in Figure 5. The location of
Figure 2d is shown in Figure 1. All maps are in a cylindrical projection. The topography in Figures 2a, 2b, and 2d is from
MOLA [Smith et al., 2001] shown with a resolution of 128 pixels/degree, and the topography in Figure 2c is from the
National Elevation Data Set (NED) for Earth [Gesch, 2007; Gesch et al., 2002] relative to the NAD83 datum. The latitude and
longitude range of each panel is as follows: (Figure 2a) 10–39°S, 210–240°E; (Figure 2b) 12–19°S, 212–218°E; (Figure 2c)
38.0–41.5°N, 113–118°W; and (Figure 2d) 0–4°N, 109–113°E.
compressional ridges formed as a result of a buckling instability related to the development of Tharsis
[Schultz and Tanaka, 1994]. If the lithosphere weakens with deformation, it develops a localized instability
resulting in regularly spaced faults [Montési and Zuber, 2003a, 2003b, 2003c]. One possible analog for the
buckling instability is the Central Indian Basin [Montési and Zuber, 2003a], where the tectonic deformation
is expressed in two different scales: folds in the basement with a wavelength of 200 km and an amplitude
of 2 km formed as a result of a buckling instability and closely spaced (7–11 km) faults with modest
(<1 km) offsets formed as a result of a localization instability. Although the Central Indian Basin is a clear
example of the effects of buckling and localization instabilities, the closely spaced and low relief stepwise offsets bear no resemblance to the high relief ridges in the STRB, and we do not consider it further as an analog
in this particular case. In section 3, we argue that the western STRB ridges are large-scale fault-bounded
blocks, unlike the periodic undulations and more subtle faults observed in the Central Indian Basin.
Wrinkle ridges provide an example of tectonic structures controlled by localization instabilities [Montési
and Zuber, 2003b]. However, wrinkle ridges are small-scale structures formed as a result of surface folding
above blind thrust faults [e.g., Schultz, 2000] and bare no resemblance to the much larger fault-bounded
STRB ridges.
For a possible compressional analog for the STRB, we consider the well-defined Martian lobate scarp,
Amenthes Rupes (Figures 1 and 2d), a 380 km long thrust fault scarp with a maximum relief of ~1.2 km
[Watters, 2003], located south of the dichotomy boundary to the east of the Isidis basin. Lobate scarps are
compressional structures interpreted as surface-breaking thrust faults [Watters and Robinson, 1997].
Amenthes Rupes is thought to be underlain by a thrust fault with a dip of 25–30° extending to a depth of
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
918
Journal of Geophysical Research: Planets
10.1002/2015JE004936
25–30 km [Schultz and Watters, 2001]. However, unlike the STRB ridges, lobate scarps on Mars do not typically
occur in populations with regular spacing between the ridges. Nevertheless, we take Amenthes Rupes as a
compressional analog for the STRB, representing a large-scale structure formed by a surface-breaking thrust
fault. Similar thrust faults combined with a periodic buckling instability might generate a population of
uniformly spaced compressional ridges as found in the STRB mapped and interpreted by Schultz and
Tanaka [1994], though we favor extensional strain to explain the STRB as detailed below.
2.2. Extensional Analog: Basin and Range
We suggest that the geomorphology and morphometry of the ridges of the STRB (Figures 1, 2a and 2b)
resemble the extensional Basin and Range Province in the western United States (Figure 2c). The Basin and
Range is characterized by a pattern of alternating basins and ranges in a zone of distributed extensional
deformation, which differs markedly from narrow continental rifts and is commonly described as a “wide rift”
[Buck, 1991]. Although the extension across the Basin and Range on the Earth is likely linked to plate tectonic
processes associated with the plate boundary stresses on various interactions between the Pacific or Farallon
and North American plates [Atwater, 1970; Severinghaus and Atwater, 1990; Atwater and Stock, 1998; Choi and
Gurnis, 2003; Liu and Stegman, 2011] or buoyancy forces arising due to the overthickened crust [Fleitout and
Froidevaux, 1983; Coney and Harms, 1984; Glazner and Bartley, 1984; Wernicke et al., 1987; Axen et al., 1993;
Jones et al., 1996; Sonder and Jones, 1999], similar structures could form on Mars in the absence of plate
tectonics if the appropriate conditions of strain rate, crustal thickness, heat flow, and rheology are met
(see section 5).
The total crustal extension estimates for Basin and Range vary between 10 and 300%, with most recent work
favoring values at the high end of this range [Hamilton and Myers, 1966; Parsons, 1995; Zoback et al., 1981;
Wernicke et al., 1988; Wernicke, 1992; Sonder and Jones, 1999; McQuarrie and Wernicke, 2005]. The direction
and amount of the extension and strain in the Basin and Range have varied with time and location. Strain
is not uniformly distributed, with some regions displaying extreme extension (e.g., 250% in the Central
Basin and Range) and some displaying minor extension (e.g., <10%) [Sonder and Jones, 1999; Parsons,
1995]. Strain rates, like strains, are strongly variable both spatially and temporally and are estimated to range
from 2 × 1016 to 5 × 1015 s1 [Sonder and Jones, 1999]. Global Positioning System (GPS) studies reveal that
the northern Basin and Range Province accommodates almost 25% of the ~50 mm/yr of relative horizontal
motion between the Pacific and North America plates [e.g., Dixon et al., 2000; Bennett et al., 2003], while
the eastern ~350 km wide Basin and Range region experiences only ~3 mm/yr of extension [Bennett et al.,
1999, 2003, 2007; Friedrich et al., 2003; Niemi et al., 2004; Velasco et al., 2010].
The regular ridge spacing in the Basin and Range is thought to be primarily determined by the thickness of
the strong layer [Fletcher and Hallet, 1983]. The regular spacing of tilts (distances between synformal and antiformal boundaries) [Stewart, 1980] and range domains of the Basin and Range may arise from necking
instabilities resulting from strength stratification of the lithosphere [Zuber et al., 1986]. Buoyancy forces resulting from the thickened crust are thought to be an important driver of the extension across the Basin and
Range [Sonder et al., 1987]. The style of extension has been attributed to the balance between localizing
and delocalizing effects, with wide rifting occurring when the delocalizing effects become larger than the
localizing effects [Buck, 2007], as will be discussed further in section 5 in the discussion of the tectonic and
geodynamic implications of possible wide rifting at the STRB.
The Basin and Range is a result of a complicated geometry of extensional faulting. The ranges are spaced at
an average of 25–35 km, separated by 10–20 km wide basins [Parsons, 1995]. Models of Basin and Range
structure range from horst and graben, tilted block, and listric fault models [Zoback et al., 1981]. Regional tilt
patterns are compatible with the tilted block and listric fault models but not with a simple horst and graben
model [Stewart, 1980]. Many ranges in the Basin and Range are dominated by block faulting [Hamilton and
Myers, 1966]. The faults may be listric, flattening out at a depth of less than 10 km [Moore, 1960; Hamblin,
1965; Hamilton and Myers, 1966; Mohapatra and Johnson, 1998]. Zoback et al. [1981] concluded that the
Basin and Range Province probably includes many types of faults, including listric faults that are steep at
the surface but become subhorizontal either at 4–5 km depth or at the base of the seismogenic crust
(~15 km). The faults can extend through the brittle-ductile transition zone and transition into low-angle
detachment faults [Constenius, 1996; Velasco et al., 2010]. However, earthquake focal mechanisms show that
much of present-day seismicity of the Basin and Range occurs within 15 km of the surface on high-angle
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
919
Journal of Geophysical Research: Planets
10.1002/2015JE004936
faults with dips of 45°–60° [e.g., Doser and Smith, 1989; Jackson and White, 1989; Ellis et al., 1999]. Thus, the listric
faults in the Basin and Range may simply be rotated, high-angle normal faults [Buck, 1988; Wernicke and Axen,
1988; Hamilton, 2005]. However, the large amount of extension may require accommodation by widespread
imbricate normal-fault blocks and subjacent very low angle normal faults [Wernicke and Burchfiel, 1982].
The interaction between faulting and surface processes such as fluvial erosion and mass wasting has produced a number of distinctive morphologies in the Basin and Range Province [Ellis et al., 1999]. Some of these
ranges show clear asymmetry, whereas the others are more symmetric, including cases both with displacement on a single normal fault and with a symmetrical horst bound by a pair of normal faults. The symmetric
and asymmetric ranges of the Basin and Range Province, which we consider as a good extensional analog for
the STRB, will be examined further in the topographic comparison in the next section.
A variety of tectonic features in the Solar System have been interpreted as having formed through similar
distributed extension. Grooved terrains on icy satellites, as exemplified by the Ganymede grooved terrain,
have been interpreted as a form of wide-style rifting [Dombard and McKinnon, 2001; Bland and Showman,
2007; Bland et al., 2010; Nimmo, 2004]. The Ganymede grooved terrain consists of remarkably uniform sets
of grooves and ridges with a characteristic spacing of 3–17 km [Grimm and Squyres, 1985]. Small-scale
tectonic ribbons in the tesserae terrain of Venus have also been interpreted as arising from wide-style rifting,
with a characteristic spacing of 2–6 km [Hansen, 1998; Ghent and Hansen, 1999; Ghent, 2002]. These examples
of wide-style rifting reveal a wide range in scales and morphologies.
3. Topographic Analysis
3.1. Topographic Analysis
We now use topographic profiles to provide a quantitative comparison of the South Tharsis Ridge Belt to the
compressional and extensional analogs described above. The topographic analysis was performed using
Mars Orbiter Laser Altimeter (MOLA) topography data for Mars [Smith et al., 2001], together with the
National Elevation Data Set (NED) for Earth [Gesch, 2007; Gesch et al., 2002]. We analyzed the western STRB
(Figures 2a and 2b), which contains six well-defined ridges, with typical length of 250 km, width of 60 km,
height of 1.5 km, and separations between ridges of up to 260 km. For comparison, we examined topographic
profiles of Basin and Range ridges and the Amenthes Rupes lobate scarp as extensional and compressional
analogs, respectively (Figures 2c and 2d).
Multiple topographic profiles were taken perpendicular to each ridge axis, approximately equally spaced but
avoiding areas that have clearly been modified by cratering or other processes. Each profile was examined for
the presence of steep scarps, the total relief, and the typical slopes on either side of the ridge. The mean and
±1 standard deviation range of the slope on each limb, total relief, and width of the ridges was quantified. For
the STRB, we also took one longer profile along a path that intersected multiple ridges and calculated the
mean ridge spacing. This profile was then truncated at the crest of the ridges on either end, and the power
spectrum of the Fourier series was used to identify the periodicity. A similar analysis was performed for a
single profile across an equal number of ridges in the Basin and Range.
3.2. Ridge Shape
The western STRB ridges are characterized by a range of morphologies. Some ridges exhibit symmetric
slopes, with an average value of 0.13–0.14 (note that slopes are expressed in unitless gradients; Table 1 and
Figure 3a, Profile A-A′), while other ridges show asymmetries in slope (Figure 3b, Profile B-B′). Each of the
three ridges in Figures 3a–3c show an elevation difference from one side of the ridge to the other as a result
of embayment from the east by Tharsis lava flows with impact crater retention ages ranging from Hesperian
to Amazonian [Tanaka et al., 2014; Anderson et al., 2012]. Another dominant feature of the profiles is the flattened tops of many ridges. Taken together, the steep symmetric slopes and flat tops on some ridges suggest
that they are bound by symmetric faults on either side, similar to the horst structures inferred for some Basin
and Range ridges. The asymmetric profiles of other western STRB ridges are consistent with these ridges each
being bound by a single fault, as observed in the majority of the Basin and Range ridges. Some ridges within
the western STRB show crested tops (Figure 3c, Profile C-C′) as is the case for most of the Basin and Range
ridges. These crested tops in both cases are likely the result of erosion, with the former likely occurring early
in Mars’s history when the erosion rates were reportedly higher [Golombek et al., 2006], consistent with the
highly degraded Martian valleys that mark the slopes in places.
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
920
Journal of Geophysical Research: Planets
10.1002/2015JE004936
Table 1. Topographic Measurements of Ridges in the Western South Tharsis Ridge Belt (111 Measurements of 5 Ridges), Basin and Range Province (BR; 394
a
Measurements of 15 Ridges), and Amenthes Rupes (AR; 6 Measurements)
Ridge
Average Slope West
Average Slope East
Average Height (km)
Average Width (km)
Number of Ridges
Number of Profiles
STRB
BR
AR
0.13 ± 0.04
0.13 ± 0.02
0.11 ± 0.02
0.14 ± 0.03
0.13 ± 0.02
0.01 ± 0.002
1.5 ± 0.4
0.9 ± 0.2
1.1 ± 0.1
59.0 ± 15.9
34.8 ± 7.9
71.7 ± 2.9
5
13
1
111
394
6
a
Locations of the profiles are shown in Figure S1 in the supporting information.
The slopes of the crested top ridges of the Basin and Range are mostly symmetric with average values of 0.13
(see Table 1 and Figure 3d, Profile X-X′). This large-scale topographic symmetry is in contrast with the asymmetric
tectonic geometry inferred for many of the ranges, likely as a result of both tectonic and erosional modification.
The steep scarp slopes for both the Basin and Range and the western STRB are not representative of the fault
dips but rather the combination of faulting, mass wasting, and erosion [Ellis et al., 1999]. The signatures of the
topographic profiles representative of the Basin and Range are influenced by tectonism (including interactions
between faults underlying adjacent ranges), as well as fluvial modification of the flanks of the ranges and
the primary drainages that follow the elongated basins. Nevertheless, some ridges have markedly asymmetric
profiles, in keeping with the asymmetric tectonic geometry (Figure 3e, Profile Y-Y′).
In contrast, topographic profiles of Amenthes Rupes (Figure 3f, Profile Z-Z′) differ significantly from the profiles
of both the western STRB and Basin and Range. The profiles clearly reflect the basic morphologic elements of
lobate scarps with relatively steep sloping scarp faces and gentle back slopes [Schultz and Watters, 2001;
Watters, 2003]. Average values for the scarp and back slopes are 0.13 and 0.04, respectively, which shows a
strong asymmetry with only the steeper slope corresponding to a fault scarp. Although some STRB profiles
Figure 3. Representative profiles of ridges in the (a–c) western South Tharsis Ridge Belt, (d, e) Basin and Range, and (f) Amenthes Rupes, with transect locations
shown in Figure 2. The vertical exaggeration is 40 times in each panel. Figures 3d and 3e are at one third the scale of the other panels to best show the shapes
of the ridge profiles.
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
921
Journal of Geophysical Research: Planets
a
b
c
d
10.1002/2015JE004936
Figure 4. (a) East-west profile across the STRB taken at 23°S, crossing five ridges, compared with a profile across five ridges of the Basin and Range ((b) profile extends
from 39.21°N 117.62°W to 38.62°N 116.15°W). The Fourier power spectra of the profiles show peaks at (c) 149 km and (d) 31.9 km, respectively. Locations of the
profiles are shown in Figure 2 as 1-1′ and 2-2′.
(e.g., Profile C-C′) do bear some resemblance to the Amenthes Rupes profile, the majority of the ridge profiles do
not and the average slopes and aspect ratios are distinctly different. We cannot rule out the possibility that
thrust faulting could produce symmetric ridges similar to those in the STRB, but clear examples of such on
Mars, the Moon, or Mercury are lacking.
From a purely topographic standpoint, the combination of symmetric and asymmetric ridges in the western
STRB is most consistent with the Basin and Range. Although some ridges in the STRB are crested similar to
those in the Basin and Range, the common observation of flat-topped ridges in the STRB may be a result
of the lower erosion rates that have persisted for most of Mars’ history [Golombek et al., 2006], coupled with
a more common occurrence of symmetric bounding faults in the STRB. The scarp slopes of the STRB and the
Basin and Range are similar, while the STRB differs from Amenthes Rupes in the much lower slope of the back
slope in the latter. Among the three tectonic cases, the western STRB has the highest average ridge height
and width of 1.5 and 60 km, respectively (Table 1). The Basin and Range ridges and Amenthes Rupes have
similar average heights, whereas the width of Amenthes Rupes is twice the average width of Basin and
Range ridges, reflecting the lower dip angle of the Amenthes Rupes back slope. The western STRB ridges
are somewhat greater in both height and width relative to the Basin and Range, resulting in an identical
height to width ratio of 0.025. In contrast, Amenthes Rupes exhibits a lower height to width ratio of 0.015.
Thus, from a simple topographic standpoint, the symmetric slopes, high relief, and narrow widths of western
STRB ridges are more consistent with an origin through Basin and Range-style extensional tectonism than a
compressional regime as originally proposed by Schultz and Tanaka [1994].
3.3. Ridge Spacing
Regular periodic basins and ranges are characteristic of wide-style rifting. Although the ridges in the STRB are
not nearly as extensive as those in the Basin and Range, a simple east-west profile through the central portion
of the STRB at 23°S crosses five distinct ridges with spacings between 130 and 260 km, with an average spacing of 183 km and a peak in the Fourier power spectrum at a wavelength of 149 km (Figure 4). This spacing is
much greater than the 25–35 km spacing observed in the Basin and Range [Fletcher and Hallet, 1983; Parsons,
1995] and is the most significant difference between the two provinces. For comparison, a topographic
profile was taken across five ridges in the Basin and Range. The mean ridge spacing for these ridges was
found to be 31.6 km, with spacings between 23.5 and 41.0 km. The characteristic ridge spacing from the
wavelength of the peak in the Fourier spectrum was 31.9 km, in agreement with the wavelength from longer
profiles of 31 km found by Fletcher and Hallet [1983].
The spacing of ridges or basins in a zone of distributed extension can be used to place constraints on the
thickness of an upper strong layer overlying a weak layer. This concept was first applied to the Basin and
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
922
Journal of Geophysical Research: Planets
10.1002/2015JE004936
Range, with a simple linear analysis predicting a ratio between the wavelength of the necking instability and
the thickness of the strong layer of 3.4–4 [Fletcher and Hallet, 1983]. Applying this ratio to the Basin and
Range, the characteristic spacing of ranges of 31 km predicts a strong layer thickness of 8–9 km, compatible
with evidence for a brittle-ductile transition at 10–15 km depth [Fletcher and Hallet, 1983].
Grooved terrain on icy satellites, as exemplified by the Ganymede grooved terrain, has also been interpreted
as a form of wide-style rifting [Fink and Fletcher, 1981; Nimmo, 2004]. The groove spacing of 3–17 km as
viewed at Voyager resolution implies a brittle layer thickness of 1–5 km [Grimm and Squyres, 1985], consistent
with the weak rheology of ice and high heat flow on that body. Detailed numerical modeling of the formation of
the grooved terrain revealed that the ratio between the wavelength of the instability and the strong layer thickness depends on the strain rate and rheology and can range between ~1 and 10 [Dombard and McKinnon, 2001;
Bland and Showman, 2007; Bland et al., 2010]. Similarly, the ribbon spacing on Venus of 2–6 km implies a brittle
layer thickness of 0.6–2.9 km [Hansen, 1998; Ghent and Hansen, 1999; Ghent, 2002], likely arising from a combination of the high surface temperature, elevated heat flow, and possibly a weaker felsic mineralogy in the tesserae
[Hashimoto et al., 2009; Mueller et al., 2009].
In this context, the STRB can be understood as part of a continuum of styles of wide rifting. The characteristic
spacing of the ranges of 149 km implies a lithosphere thickness of 37–50 km, consistent with the range of lithosphere thicknesses inferred for Noachian Mars [Schultz and Watters, 2001; McGovern et al., 2004; Grott et al., 2005;
Hoogenboom and Smrekar, 2006; Grott and Breuer, 2008]. The thicker lithosphere controlling STRB tectonism in
comparison to the Basin and Range is easily understood as an outcome of both the stronger diabase rheology of
the former [e.g., Mackwell et al., 1998; Grott and Breuer, 2008] and weaker quartz-dominated rheology of the latter.
We also note a progression in morphology from ridges that are remarkably periodic, uniform, and continuous at
the smallest scales (the Ganymede grooved terrain) to ridges that are both less periodic and more irregular at the
largest scales (the STRB). This progression suggests a scale dependence in which the complexity of faulting in
thicker and stronger lithospheres results in a less uniform pattern of basins and ranges.
Alternatively, periodic ridges can also arise in compressional tectonism from a localization instability. An analysis
of wrinkle ridges on Mars found that their characteristic spacing of 40 km in Solis Planum and 90 km in the
northern lowlands is consistent with lithosphere thicknesses of 35 km and 60 km, respectively [Montési and
Zuber, 2003b]. Those results imply a ratio between the dominant wavelength of the structures and the lithosphere thickness of 1.1–1.5. Applying this to the wavelength of the STRB ridges yields a lithosphere thickness
of 100–135 km, which is difficult to reconcile with estimates of the lithosphere thickness of Noachian Mars.
We also note that some ridges in the northern STRB have an apparent en echelon pattern (Figure 2), which is
unlike both the compressional and extensional analogs. This pattern may indicate a component of shear deformation in their formation that may have been coupled with either extensional or compressional tectonism.
4. Strain Indicators
4.1. Deformed Craters of the Western South Tharsis Ridge Belt
Previous work indicated that deformed craters support the compressional interpretation of the STRB [Schultz and
Tanaka, 1994], and so we now revisit the evidence for deformed craters in this region. The western STRB contains
many impact craters that may help resolve the strain history of the region. The majority of undeformed craters
observed on Mars, Venus, and the Moon in general are circular, with only ~5% of the population of small
(<150 km diameter) craters being elliptical due to oblique impacts [Bottke et al., 2000; Andrews-Hanna and
Zuber, 2010]. Systematic departures from circularity indicate significant deformation and may be used as indicators of the magnitude and, in some cases, sense of the strain [Tanaka and Golombek, 1994; Pappalardo and Collins,
2005]. Extensional and compressional mechanisms of the western STRB formation may be distinguished based
on the observation of craters that are elongated or shortened in the direction orthogonal to the ridge axes,
respectively. However, although crater analysis can in some cases reveal the strain history, it can be nontrivial
in regions where strain is highly variable both locally and regionally, as may be the case in the western STRB.
Part of the crater population of the western STRB has been destroyed by magmatic, fluvial, aeolian, and
colluvial processes [Anderson et al., 2012]. Nevertheless, the crater retention ages of the ranges and basins
are Middle Noachian with the ranges yielding an older age, as expected from the subsequent partial infilling
of the basins. When considering the destruction of part of the impact crater population, the materials which
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
923
Journal of Geophysical Research: Planets
10.1002/2015JE004936
Figure 5. MOLA topographic maps in cylindrical projection detailing impact craters in the western STRB (measurements of
the crater dimensions used true distance and are not affected by the image projection). Capital letters denote the crater
names used in the text. See Figure 2a for the locations of craters.
compose the ranges and the basement of the basins are likely Early Noachian in age, consistent with highstanding materials to the west [Anderson et al., 2012]. The younger age of the ranges and basins, when compared
to the relatively high-standing materials to the west, is likely due to the aforementioned modification processes,
enhanced in the STRB by the structurally controlled steeper slopes.
The remaining impact craters that were not destroyed by the aforementioned processes are mostly circular to
slightly elliptical in shape, with well-preserved rims (Figure 5). After a survey of the aspect ratios of a large
number of well-preserved impact craters in the region, we find only six craters that show possible evidence
for departures from circularity (Figure 5). We classify these as either elongated or shortened in the direction
orthogonal to the ridge axes. We note that the ridges likely formed over a protracted period of time, whereas
individual craters formed instantaneously. Thus, some craters that cut into the ridges, clearly postdating
some or even much of the ridge-forming tectonism, may still be deformed by ongoing deformation.
Crater A (Figure 5) could be interpreted as shortened due to the orientation of its short axis perpendicular to the
ridge axis. However, its aspect ratio of 1.03 is not significantly different from 1 given the natural variability in crater
rim shapes and the effects of the crater degradation through time, and this crater clearly postdates much of the
ridge-forming tectonism. Crater B has a higher aspect ratio of 1.21, although it is unclear whether this crater
should be classified as shortened or elongated, as the long axis of the crater is oblique to the nearby ridge axis,
and the ridge axis appears to be rotating as it approaches the crater. The mean strike of the ridge is nearly orthogonal to the short axis of the crater, consistent with tectonic shortening of the crater in the direction orthogonal
to the ridge. The end of the ridge closest to the crater strikes at an angle of ~45° to the short axis of the crater,
consistent with either elongation or shortening. If the rotation of the ridge axis toward the south is assumed to
continue to the crater itself, then the crater would be lengthened perpendicular to this extrapolated ridge axis
orientation. The interpretation of this crater is further complicated by the partial volcanic burial of the western
rim, superposed crater on the eastern rim, and degradation of the impact crater rim since its formation.
Crater C is located in the southern part of the western STRB and is associated with an area of high topography,
displaying an aspect ratio of 1.38. Its long axis direction is perpendicular to the regional trend of the western
STRB ridge axes suggesting lengthening. This is supported by the observation that the portions of the crater
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
924
Journal of Geophysical Research: Planets
10.1002/2015JE004936
rim parallel to the long axis are discontinuous and lower in relief, suggestive of extensional necking of the rim.
Crater D is located on top of a ridge, thus likely postdating at least some of the ridge-forming tectonism, and
its aspect ratio is calculated as 1.09. The orientation of its long axis perpendicular to the ridge axis suggests
elongation, though this could result from flexural extension associated with compressional buckling of the
ridge. Finally, crater E clearly predates the superposed ridge, with an aspect ratio of 1.02. Its short axis is perpendicular to the ridge axis suggesting shortening. However, this crater is part of a chain of three craters that
appear to have formed simultaneously, and the expansion of its rim in the downrange direction may have
been affected by the simultaneous formation of the adjacent crater to the east, affecting the aspect ratio.
This region also contains a small number of highly tectonically deformed ancient basins of possible impact
origin, such as Koval’sky basin (Figure 5f) [Werner, 2008]. If impact in origin, these irregular basins have undergone substantial deformation. The linear ridges and scarps, interpreted to be highly degraded impact crater
rims and shown as dashed lines in Figure 5f, suggest tectonic deformation of Koval’sky basin. This deformation likely involved a significant component of shearing to produce the observed distorted shape. Due to the
significant tectonism and high amount of strain, as well as modification by other processes as noted above, it
is not possible to decipher the strain history of this basin. It is possible that tectonic structures in the preimpact basement crust may influence the shapes of impact craters, as observed in the lunar basins (e.g., Crisium
and Tsiolkovskiy) [Fielder, 1965; Kopal, 1966; Guest and Murray, 1969; Spudis, 1993], in Venus [Aittola et al.,
2007], Argyre region of Mars [Öhman et al., 2008], Manicouagan Crater, Quebec [Floran and Dence, 1976],
and Meteor Crater, Arizona [Shoemaker, 1963; Quaide et al., 1965]. However, it is unlikely that preexisting
structures could result in the extreme departure from circularity observed at Koval’sky.
Although this region has clearly undergone a significant amount of tectonic deformation, the aspect ratios of
impact craters do not display a clear trend of shortening or lengthening orthogonal to the ridges. As a result,
it is not possible to unambiguously distinguish between extensional and compressional formation mechanisms by this approach. The lack of deformed craters is not surprising, given the highly localized nature of the
deformation along the edges of the ranges and the extensive erosion and resurfacing since the formation of
the ranges. Craters that are small relative to the spacing between the ridges are unlikely to have experienced
tectonic elongation by the wide rifting. The observation of both elongation and shortening of the craters may
suggest that this region has undergone several different tectonic regimes with reversing senses of strain, as
has previously been suggested for Tempe Terra on Mars [Tanaka and Golombek, 1994]. This scenario is
consistent with the presence of crosscutting graben, strike-slip faults, and wrinkle ridges in this area
[Andrews-Hanna et al., 2008], as well as the observation that Hesperian-aged wrinkle ridges commonly appear
to parallel and perhaps nucleate from the more ancient STRB ridges. However, given the possibility of
elliptical craters forming through low-angle impacts or from near-simultaneous impacts in crater chains,
no firm conclusions regarding the strain history can be drawn from the crater population.
4.2. Strain From Crustal Thickness
Crustal thickness models may provide a better-preserved record of the large-scale strain across this tectonic
province. Extension is expected to result in crustal thinning, while compression should cause thickening. The
western STRB is contained within the western trough surrounding Tharsis, which is a result of the membraneflexural deformation of the lithosphere in response to the volcanic load [Phillips et al., 2001]. Separate from
this long-wavelength signature, crustal thickness models [Zuber et al., 2000; Neumann et al., 2008] indicate
that the crust comprising the western STRB is thinner than that in the highlands to the west (Figures 6a
and 6b). Although the eastern portions of the STRB overlap with the crustal thickening associated with
Tharsis which likely arose later in time, the western zone of the ridge belt exhibits a thinner crust than the
highlands to the west by ~2–5 km. The crust thins again farther to the west, in western Terra Sirenum, but this
region is not representative of typical highland crust and appears to have undergone a separate but possibly
related episode of deformation (see section 4.4). Although the reduction in crustal thickness within the western STRB is small, the width of the western STRB of up to 1000 km (corresponding to spherical harmonic
degree 11) is well resolved in the Mars Reconnaissance Orbiter gravity model [Zuber et al., 2007]. The thinned
crust of the western STRB is associated with the lower topography between the ridges, as would be expected
from isostatically compensated crustal thinning (Figures 6c and 6d).
Given that the crust in the southern highlands is remarkably uniform in thickness outside of the giant impact
basins and Tharsis [Neumann et al., 2008], the thinner crust in the STRB may be interpreted as an effect of the
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
925
Journal of Geophysical Research: Planets
10.1002/2015JE004936
Figure 6. (a) Crustal thickness map of the western STRB based on the model of Neumann et al. [2008], (b) representative
crustal thickness profile, (c) topographic map, and (d) representative topographic profile. Dashed lines indicate the typical crustal thickness and elevation outside the western STRB. The contour interval in Figure 6a is 10 km. Location of the
representative profile is shown in Figures 6a and 6c with white line. All maps are in a cylindrical projection, and the profile
follows a great circle path chosen to avoid large impact craters.
formation of this ridge belt. This thinning of the crust is consistent with extensional deformation. Assuming
that the crust in this region had an initial thickness of 60 km, similar to the crust to the west, the observed
thinning by ~2–5 km based on individual crustal thickness profiles suggests an extensional strain of 0.03
to 0.09. The horizontal extension in the region is calculated from this strain to be ~30–90 km, based on
width of the ridge belt of ~1000 km.
Thinning of the crust contradicts the compressional interpretation of the ridges and instead supports horizontal extension of the crust. Thus, both the western STRB and Basin and Range display crustal thinning
and a significant amount of extension. The total extension in the Basin and Range varies between 10 and
300% [Hamilton and Myers, 1966; Parsons, 1995; Zoback et al., 1981; Wernicke et al., 1988; Wernicke, 1992;
Sonder and Jones, 1999; McQuarrie and Wernicke, 2005], with the inferred western STRB strain approaching
the low end of that range. However, as revealed in the Basin and Range Province, crustal thickness can only
set a lower bound on the amount of extension because of the effects of substantial lower crustal flow. The
strain based on the crustal thinning will be compared with an estimate of the amount of extensional tectonic
deformation inferred from the relief of the ridges in the next section.
4.3. Strain From Horizontal Tectonic Deformation
If the deformation within the western STRB was purely tectonic in nature, the observed relief across the ridges
could be used to estimate the amount of extension or compression for assumed fault dips. The horizontal
tectonic deformation for each ridge and the province as a whole can be calculated from the displacement
across the faults bounding the ridges for both the extensional and compressional interpretations. For this
analysis, we assumed that each ridge is bordered by two symmetric faults and then calculated the horizontal
deformation generated by these inferred faults using topographic profiles. First, topographic profiles were
taken perpendicular to the ridge axes. For a single profile, the vertical throw (y) was calculated by subtracting
the typical elevation outside of the ridge from the maximum height of the profile. Then, the horizontal deformation (x) was calculated for fault dip angles α ranging from 0° to 90° as x = 2y/tan(α), assuming that each
ridge is bound by a symmetric pair of faults. The horizontal deformation from all profiles crossing each ridge
was numerically integrated to calculate the change in the horizontal area. The same procedure was carried
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
926
Journal of Geophysical Research: Planets
10.1002/2015JE004936
Figure 7. Strain inferred from the relief of the ridges as a function of the dip angle for (a) western STRB and (b) Basin and
Range. The western STRB strains are expressed as extensional (negative values), but the magnitude of the inferred strain is
also valid for compressional strain arising from thrust faults with equivalent dips.
out for each ridge, and the average strain across the province was calculated by summing the areal horizontal
deformation from each ridge and dividing by the area of the western STRB.
The areal strains calculated for normal faulting are in the range of 1.6 × 103 to 0.6 × 103 for fault dips
ranging from 45° to 70°, respectively (Figure 7a). If each ridge was bordered by a single fault, as is the case
for many Basin and Range ridges, the strain would be reduced by half. The extensional strain is underestimated by this analysis, since it does not include the effects of fault rotation, erosion of the footwalls,
and deposition on the hanging walls. However, in order to accommodate the strain of up to 0.09 inferred
from the crustal thinning, exceedingly gentle fault dips would be needed (<5°). This dip is less than the
observed average dips of the scarps bounding the ridges, which place a lower bound on the fault dip,
and is also less than the frictional lock-up angle for normal faults [Scholz, 2002]. This suggests that the
amount of strain is underestimated if only faulting is considered, and some additional mode of extension
is required. This is consistent with the Basin and Range where the total extensional strain of 0.1 to 3.0
greatly exceeds the strain that would be calculated based on ridges bounded by normal faults with typical
dips [Hamilton and Myers, 1966; Parsons, 1995; Zoback et al., 1981; Wernicke, 1992; Sonder and Jones, 1999;
McQuarrie and Wernicke, 2005]. A similar strain calculation was performed for the Basin and Range assuming that the ranges are each bordered by a single normal fault. The average near-surface fault dip angles
between 45° and 60° [e.g., Doser and Smith, 1989; Jackson and White, 1989; Ellis et al., 1999] result in tectonic
strain estimates in the range of 0.014 to 0.008, respectively (Figure 7b). This result is much less than
the inferred strain across the province, and is consistent with the interpretation that the Basin and
Range extension was accommodated by a combination of listric normal faulting and ductile deformation
[Mohapatra and Johnson, 1998]. Similarly, our analysis above suggests that a significant fraction of the
extensional strain in the western STRB must have been accommodated by ductile deformation of the crust
and/or listric normal faulting.
If the western STRB ridges are instead bounded by symmetric thrust faults, the inferred areal strain is 0.003
for the optimum thrust fault dip of 30°. This would result in thickening of the crust by 0.2 km, which could
not be resolved in the crustal thickness models relative to the background variability. Nevertheless, the
observed crustal thinning in the western STRB conflicts with the compressional interpretation.
4.4. Possible Link to Deformation in Western Terra Sirenum
If early Mars lacked plate tectonics, then the formation of the STRB by extensional tectonism must have been
accommodated by compression elsewhere. One candidate region to accommodate this extension is found
in western Terra Sirenum (Figure 1). Although the craters in the western STRB did not yield a consistent sense
and orientation of strain, ~500–700 km farther to the west is a belt of intensely deformed craters and basins
in western Terra Sirenum, with shapes that depart significantly from the expected circular or elliptical
outline (Figure 8). The irregular shapes of the seven largest impact basins in the western Terra Sirenum
region clearly suggest that this area has undergone major deformation of the crust, and the lack of observed
straight rim segments or tectonic offsets suggests that this deformation was distributed in nature, possibly
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
927
Journal of Geophysical Research: Planets
10.1002/2015JE004936
due to continuous plastic/ductile
rather than tectonic deformation
(Figure 8). These basins are some
of the largest and oldest strain markers in this region, with diameters
between ~175 and 295 km. In contrast,
craters with diameters smaller than
~170 km appear to be undeformed,
with sharp rims and well-defined
circular shapes.
In order to quantify the deformation
in western Terra Sirenum, we measured the strain of each highly
degraded crater and used rose diagrams and strain ellipses to represent
our results (Figure 9). We analyzed
204 craters with diameters greater
than 10 km. The crater rims were
traced using MOLA topography overlain with Thermal Emission Imaging
System (THEMIS) daytime infrared
data. The best fit circles to the crater
rims from a Mercator projection were
found by a least squares analysis. The
radius of the best fit circle was subFigure 8. MOLA map in Mercator projection showing craters of western
tracted from the mean radius of the
Terra Sirenum, the seven largest (1–7) having diameters >170 km. The
figure extends from 15–55°S to 175–215°E.
crater and normalized to the mean
radius to give the strain as a function
of azimuth. Positive and negative strain values represent relative compression and extension, respectively
(these relative strains are equivalent to the absolute strains if the mean crater radius was not changed during
the deformation). In order to quantify the deformation, the maximum, minimum, and root-mean-square
(RMS) strains were calculated for each crater. Rose diagrams (Figure 10) were then used to represent the
strain amount and direction relative to the best fit circle for each crater. If compressional deformation in
western Terra Sirenum accommodated the extension inferred for the STRB, these craters would be expected
to be shortened along axes pointing toward the STRB (approximately NE).
Rose diagrams of the heavily deformed basins >170 km in diameter indicate that this region experienced significant deformation, dominated by NE-SW compression and/or NW-SE extension (Figure 10b). The absolute
sense of strain cannot be determined without knowledge of the initial crater radius, and so it is not possible
to distinguish between compression in one direction and extension in the perpendicular direction. Individual
basins display different strain directions, but the average rose diagram confirms the regional trend
(Figure 10a). The axis of maximum shortening is oriented approximately in the direction of the western
STRB, consistent with the interpretation that extension in the western STRB could have been accommodated
by compression in western Terra Sirenum. The average root-mean-square (RMS), maximum (shortening), and
minimum (elongation) strains for the seven heavily deformed basins are 0.09, 0.18, and 0.17, respectively.
The RMS, maximum, and minimum horizontal deformation for these basins are 180, 360, and 340 km,
respectively. These basins are not large enough for the effects of planetary curvature to increase the likelihood of elliptical basin formation [Andrews-Hanna and Zuber, 2010]. The consistent orientation of the shortening of these basins is not compatible with their current shapes being the result of oblique impacts. For
comparison, the average rose diagram from all craters 10–170 km in diameter shows similar orientations of
shortening and elongation to the seven heavily deformed basinsbut is lower in magnitude. The RMS, maximum, and minimum strain values for the craters 10–170 km in diameter are 0.04, 0.085, and 0.091. The compressional strain inferred from the craters and basins in western Terra Sirenum is comparable in magnitude to
the strain inferred across the western STRB from the crustal thickness models (~0.03–0.09).
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
928
Journal of Geophysical Research: Planets
10.1002/2015JE004936
Figure 9. (left) MOLA topographic maps in Mercator projection detailing large deformed basins (estimated crater rims
shown by dotted lines, best fit circle by white dashed lines, and number at top left correlative with numbered impact
crater in Figure 8) in western Terra Sirenum. (right) Rose diagrams showing extensional (red) and compressional (blue)
strain values as a function of azimuth relative to the best fit circle.
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
929
Journal of Geophysical Research: Planets
10.1002/2015JE004936
Figure 10. Average strain rose diagrams for (a) the seven large, highly deformed basins shown in Figure 8 and (b) all craters
diameters 10–170 km in diameter. The orientations of the strain were binned in 10° intervals to generate the rose diagram.
However, the rose diagrams of the individual basins show highly variable strain that is not consistent with
simple shortening or elongation. Significant departures from an elliptical shape indicate significant spatial
variations in the strain on 100 km length scales. The distorted shapes of the deformed basin rims, lacking
linear rim segments or discrete tectonic offsets, may indicate ductile rather than brittle/tectonic deformation.
This differs from the morphology of the oldest impact basins in the western STRB, which display apparent
tectonic deformation with linear crater rim walls (Figure 5f). This change in deformation style suggests a
lateral transition between brittle and ductile deformation between these two regions or possibly some form
of crustal mobility that varies between the two regions.
The timing of the deformation in the western STRB is not easily constrained, since the majority of exposed
craters appear to have escaped modification due to the localized nature of the deformation. Furthermore,
part of the crater population has been destroyed by volcanic-, wind-, and water-related resurfacing as noted
above [Anderson et al., 2012]. The interpretation of the history of deformation may be complicated by the
interaction of the western STRB with major stages of Tharsis development, including reactivation of basement structures [Anderson et al., 2012]. It is also possible that the formation of individual ridges or groups
of ridges may have occurred at different times and by different processes. However, the distributed deformation in western Terra Sirenum resulted in a large population of modified craters, enabling us to constrain the
timing of the deformation there using crater size-frequency distributions [Hartmann and Neukum, 2001]. The
largest heavily deformed craters fall on an isochron together with a subset of the smaller craters (>90 km in
diameter), with a crater retention age of 4.02 ± 0.03 Ga, placing them in the Early Noachian (Figure 11). This is
comparable to the crater retention age of Middle Noachian determined for the basins and ranges of western
STRB and Early Noachian for the highland materials directly west of the western STRB in northern Terra
Sirenum [Anderson et al., 2012]. We emphasize that these are model ages only and that significant uncertainties exist in any age from crater counting, particularly those at the time of the late heavy bombardment of the
inner Solar System [Gomes et al., 2005]. If part of the crater population of this region was destroyed by this
deformation or by other processes, the crater retention age would underestimate the true age of the major
deformation. Less deformed smaller craters with diameters of 30–60 km yield an age of 3:88þ0:01
0:02 Ga, indicating
that deformation had largely ceased by that time. Thus, intense deformation of western Terra Sirenum likely
began sometime during the Early Noachian or pre-Noachian period and continued at least into the Middle
Noachian and possibly into the Late Noachian [Dohm et al., 2013]. Possible volcanism or other resurfacing
processes led to the removal of a significant fraction of the small (diameters <30 km) undeformed craters
after ~3.7 Ga, as supported by a rollover in the crater size-frequency distribution at that diameter. This is consistent with other regions of Mars such as the Argyre basin with crater populations pointing to major resurfacing during the Late Hesperian/Early Amazonian interpreted to correlate with major Tharsis-driven
environmental change, possibly at global scale [Dohm et al., 2015, and references therein]. If the deformation
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
930
Journal of Geophysical Research: Planets
10.1002/2015JE004936
in western Terra Sirenum was responsible for accommodating the strain in
the western STRB, it would have been
synchronous with western STRB development, implying a similar history for
STRB tectonism. However, we emphasize that a direct link between deformation in western Terra Sirenum and the
STRB is speculative at this point. Other
ancient geologic provinces of Mars also
show evidence for dynamic activity
early in Mars’s history [Dohm et al.,
2013, 2015].
5. Rift Modeling
5.1. Controls on the Style of Rifting
Although we cannot rule out a compressional origin for the western STRB,
multiple lines of evidence support the
extensional interpretation. Here we
examine the geodynamic implications
of Basin and Range-style extension on
Figure 11. Crater size-frequency distributions and isochron fits for the Mars. Extension in the Basin and Range
craters in western Terra Sirenum. Dark grey squares denote the craters
on Earth is thought to be driven largely
with diameters 30–60 km, and black squares denote craters with diameters by plate tectonic processes associated
>90 km.
with the subducting Farallon Slab [e.g.,
Liu and Stegman, 2011], though a mantle
plume [Parsons et al., 1994], plate boundary stresses [e.g., Choi and Gurnis, 2003] and buoyancy forces [e.g., Sonder
and Jones, 1999] may also play roles. However, in the context of our analogy to Mars, the immediate factors governing the extension can be considered independent of the broader geodynamic setting. The style of extensional
deformation depends on the heat flow, strain rate, crustal thickness, and rheology of the crust and mantle. On
Earth, lithospheric extension can take the form of a narrow rift (e.g., the East African rift valley [Kogan et al.,
2012]), a wide rift (e.g., the Basin and Range Province of the western United States [Dickinson, 2002]), or a core
complex (e.g., southern Tibet [Saylor et al., 2010]). The observed mode of lithospheric extension places constraints
on those parameters. We have applied a model of pure shear lithospheric extension [Buck, 1991] to the western
STRB in order to predict the mode of rifting for early Mars under a range of conditions in an effort to constrain
the parameters that may have led to wide-style rifting in the western STRB. Previous studies have applied
this approach to extensional deformation of other regions within Tharsis, including Valles Marineris [Anderson
and Grimm, 1998], Tempe Terra [Wilkins et al., 2002], Alba Patera [Kiefer and Johnson, 1995], and Coracis Fossae
[Grott, 2005], as well as the icy satellites [Nimmo, 2004].
Lithospheric extension is controlled by the balance between the yield stress, thermal buoyancy, and compositional buoyancy [Buck, 1991]. The mode of lithospheric extension is determined by the localization or delocalization of the strain (i.e., extension may stay in weak regions or migrate to other places, depending on the evolution
of the strength of the lithosphere [Buck, 2007]. Localized deformation is a consequence of weakening of the lithosphere in response to the extension and results in narrow rifting. Lithospheric extension and thinning result in
advective heat transfer, increasing the temperature gradient. The elevation of the isotherms reduces the strength
of the crust and mantle since both have a temperature-dependent rheology. This will localize the strain and stress
within the thinned lithosphere, leading to either a narrow rift or ductile necking of the crust [Buck, 2007]. Necking
is a self-reinforcing process since the extension becomes more rapid when the necked area is weaker. Buoyancy
effects also contribute to localized rifting. The flow of hot, low-viscosity lower crust into the region of crustal
stretching reduces the crustal thickness variations. This crustal flow enhances the crustal buoyancy beneath
the rift, which contributes to the extension. Thermal buoyancy also contributes, with fast extension leading to
an elevated thermal gradient and enhanced uplift, which in turn drives further extension.
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
931
Journal of Geophysical Research: Planets
10.1002/2015JE004936
Table 2. Model Parameters
Wide rifting occurs when the delocalizing effects are greater than the
2
localizing effects. Thermal diffusion
3.72
g
Acceleration of gravity
m/s
2
6
has a delocalizing effect and leads
10
κ
Thermal diffusivity
m /s
3
ρc
Crustal density
kg/m
2700
to thickening and strengthening of
3
Mantle density
kg/m
3500
ρm
the brittle part of the lithosphere,
1
5
α
Thermal expansion coefficient
K
3 × 10
which favors wide rifting. Extension
also replaces crust with stronger
mantle material, which has a further delocalizing effect. The resultant increase in the strength of the rift
causes the locus of extension to shift to adjacent undeformed areas, leading to wide rifting.
Symbol
Name
Units
Value
The strain rate, crustal thickness, and heat flow are the main factors influencing the style of extension in this
model. The strain rate influences the style of extensional deformation both because it affects the balance
between lithospheric thinning and thermal diffusion and because the viscous stresses in the crust and mantle
depend strongly on the strain rate. The mean strain rates on Mars are poorly constrained but are thought to
be considerably lower than those in terrestrial rifts such as the Basin and Range (2 × 1016 to 5 × 1015 s1)
[Sonder and Jones, 1999]. In our computations, the strain rates are varied between 1014 and 1017 s1,
encompassing the range of strain rates in the Basin and Range Province as well as typical values assumed
for early Mars [e.g., McGovern et al., 2004; Grott and Breuer, 2010; Ruiz, 2014].
The crustal thickness and its variation with time also play an important role in the nature of extensional
deformation of the lithosphere. The average thickness of Martian crust has been calculated from gravity
and topography data to be between 38 and 62 km [Wieczorek and Zuber, 2004]. The mean crustal thickness
within the western STRB is 55 km, but this has likely been thinned relative to the mean value of 60 km to the west.
The rheology of the crust and mantle affects the strength of the lithosphere and therefore the mode of rifting. For
Earth, dry and wet quartz and pyroxene rheologies are used to represent the crust. For dry quartz and pyroxene,
all extensional modes (core complexes and both wide and narrow rifting) are predicted on Earth, whereas for wet
quartz only narrow mode and core complexes are predicted [Buck, 1991]. Both wet and dry diabase rheologies
have been adopted for the Martian crust [Grott et al., 2005; Grott and Breuer, 2008]. In our computations we have
used wet diabase [Caristan, 1982] for the crust and wet olivine [Karato and Wu, 1993] for the mantle. The thermal
gradient also affects the mode of rifting both by its control over the rheology and by direct thermal buoyancy
effects. Estimates of the thermal gradient during the Noachian range from 15–20 K/km [Schultz and Watters,
2001] to >20 K/km [McGovern et al., 2002, 2004] to 27–33 K/km [Grott et al., 2005].
5.2. Methodology
In order to investigate the thermal effects of rifting, the total change in force balance during extension is calculated
as the sum of the effects of thermal buoyancy, crustal buoyancy, and yield strength. The sign of the change in the
total force balance during extension determines whether extension is localized (negative force balance) or
delocalized (positive force balance). A negative change in force balance at low thermal gradients leads to localized
extension in a narrow rift, with the lower temperatures causing the brittle/tectonic deformation in the rift. A
positive change in force balance leads to delocalized extension in a wide rift or a distributed zone of deformation.
A negative change in force balance at higher thermal gradients leads to localized extension in a core complex,
with the higher temperatures causing the ductile deformation in the core complex.
Temperature changes were modeled at the center of the extending region using the one-dimensional
advection-diffusion equation [Buck, 1991; Grott, 2005]:
∂T
∂2 T
∂T
¼ κ 2 vz ;
∂t
∂z
∂z
(1)
where T is temperature, t is time, vz is the downward vertical velocity (vz = έz, where έ is the strain rate and z
is depth), and κ is the thermal diffusivity. The thermal conductivity of the Martian crust and mantle is thought
to vary between 2.5 and 3.5 W m1 K1 [Grott and Breuer, 2008]. Crustal densities for the southern highlands
have been estimated to be 2500–3000 kg m3 [McGovern et al., 2004], leading to an estimated thermal
diffusivity of ~1 × 106 m2 s1. The parameters used in the calculations are given in Table 2. In our model
the lithosphere was extended to a total accumulated strain of 2.5% at strain rates of 1017–1014 s1,
corresponding to rifting velocities (vx) of 0.03–3 cm/yr and time spans (δt) of 80–0.8 Myr, respectively.
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
932
Journal of Geophysical Research: Planets
10.1002/2015JE004936
The initial temperature was calculated from the steady state solution with a constant surface temperature (Ts)
of 220 K for a given heat flux qm at the base of the mantle and a thermal conductivity of 2.9 W m1 K1. As in
previous studies [Buck, 1991; Grott, 2005], we have neglected radiogenic heat production within the crust, but
this is expected to have a small effect on the results.
The advection-diffusion equation (1) was solved using an implicit finite difference scheme, with a mesh size
of Δz = 1 km in the vertical direction. Boundary conditions were determined by fixing the temperature at the
surface and by specifying constant heat flow at the base of the crust. The change in the thermal buoyancy
(the density anomaly due to thermal expansion) is given by Grott [2005] and Buck [1991]
dF tb ¼ g∫0 ραΔT ðzÞzdz;
L
(2)
where L is the lithospheric thickness, α is the thermal expansion coefficient, ρ is the density, and ΔT denotes
the temperature change at the center of the rift during the extension. For simplicity, we set ρ = ( ρc + ρm)/2 in
calculating the thermal buoyancy, where ρc and ρm are the crust and mantle densities, respectively. The
thermal buoyancy change in equation (2) corresponds to the change in pressure due to thermal buoyancy
integrated through the vertical column or the equivalent horizontal force per unit length parallel to the rift
and is not strictly speaking a force. The base of the lithosphere is defined as the brittle-ductile transition depth.
The change in crustal thickness h during the extension is described by Buck [1991], Nimmo [2004], and Grott [2005]
∂h
∂2 h
∂h
∂u
¼ κf 2 v x h ;
∂t
∂x
∂x
∂x
(3)
where κf is the effective flow diffusivity and vx is the applied horizontal rifting velocity. For a zone of extension
between horizontal distances of Xe and Xe, the rifting velocity vx varies linearly between zero at the rift
center (x = 0) and ±έXe at the ends of extension zone at x = (±Xe). Equation (3) is discretized and solved using
an implicit finite difference scheme between x = 0 and XL, where XL is the half width of a region of initially
undeformed lithosphere (the full model domain). The boundary conditions are ∂h/∂x = 0 at x = 0 and XL.
The half widths of the regions of extension and the full model domain are taken to be Xe = 40 km and
XL = 250 km, respectively. The effective flow diffusivity kf depends on time and is given by Buck [1991]
3
gΔρ
hRT 2M
E
;
(4)
exp
κf ¼
RT M
C E ðT M T S Þ
where C = A1 106 Pa1 n is the preexponential term in the effective viscosity relation, A is a material-dependent
constant, Δρ* = (ρm ρc)ρc/ρm, where TM is the Moho temperature. The change in crustal buoyancy force is
computed by
h
ρ
dF ch ¼ ∫0 Δpdz ≈ ghΔhð ρm ρc Þ c ;
(5)
ρm
where Δh denotes the crustal thickness change at the middle of the rift during the extension. As with the
other forces described in this section, this is not strictly a force but rather a pressure integrated vertically
through the lithospheric column or a force per unit length parallel to the rift.
The vertically integrated yield strength of the crust is given by
F ys ¼ ∫0 σ ðT; z; ε_ Þdz;
h
(6)
where σ is the lesser of the frictional strength or the ductile strength. Following Grott and Breuer [2008], the
frictional stress σB for extension is calculated according to Byerlee’s law by neglecting the yield stress and
only considering the frictional strength on already-failed planes
σ v ≤ 529:9 MPa
0:786σ v ;
σb ¼
;
(7)
56:7 MPa þ 0:679σ v ; σ v > 529:9 MPa
where σv = βρrgz is the lithostatic pressure [Mueller and Phillips, 1995], ρr is the density of the overlaying rock,
and β is the factor reducing the lithostatic pressure due to the effects of pore fluid pressure and regional stresses. Given extensive evidence for higher rates of erosion and active surface hydrology during the Noachian,
hydrostatic pore pressures in the crust are the most likely scenario. Therefore, we have scaled the lithostatic
pressure with a factor of β = 0.6 to approximate the reduction in stress due to both horizontal tension and
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
933
Journal of Geophysical Research: Planets
10.1002/2015JE004936
15 1
16 1
Figure 12. Rifting mode diagrams for strain rates of (a) 8 × 10
s and (b) 1 × 10
s . (c) Total forces for a 60 km
15 1
thick crust and a strain rate of 5 × 10
s and (d) rifting mode diagram for a 60 km thick crust. The range of thermal
gradients estimated for the Noachian period is indicated by the shading in Figure 12c and the range of strain rates estimated
16
15 1
for the Basin and Range (2 × 10
to 5 × 10
s ) is indicated by the shading in Figure 12d.
hydrostatic pore pressure [Kohlstedt et al., 1995]. We also investigated the effect of pore fluid pressure ratio (β) for
the computation of the brittle stress and found out that the yield strength change is negligible (<0.03 TN/m) for
the thermal gradients of 10–50 K/km.
The ductile flow stress (the stress to cause ductile deformation at a specific strain rate and temperature) is
1=n
E
ε_
;
(8)
σd ¼
exp
nRT
A
where T is the absolute temperature, R is the universal gas constant, and E is activation energy. The parameters
adopted for the calculations are given in Table 3.
5.3. Results
We computed the change in the total force balance from equations (4), (6), and (7) for a range of thermal
gradients, crustal thicknesses, and strain rates (Figure 12). We first consider the predicted mode of extension
Table 3. Rheological Parameters Assumed for Mars
n
a
A
n 1
Material
(Pa
b
Wet diabase
c
Wet olivine
3.05
3.0
a
A is a material dependent
b
Caristan [1982].
c
1
E
s
)
C=A
1
(KJ mol
20
3.1 × 10
15
1.9 × 10
276
420
)
6
1n
(1.0 × 10 Pa)
(Pa s)
7
1.6 × 10
526.32
constant, E is the activation energy, and n is a material-dependent exponent.
Karato et al. [1986].
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
934
Journal of Geophysical Research: Planets
10.1002/2015JE004936
for two representative strain rates, 8 × 1015 and 1 × 1016 s1 (Figures 12a and 12b). We find that wide
rifting is favored for a thin crust at high thermal gradients or for a thick crust at lower thermal gradients.
Our model results reveal that for lower strain rates and corresponding smaller rifting velocities, wide rifts
occur for a larger range of thermal gradients and crustal thickness (Figures 12a and 12b).
We now examine the implications of wide rifting at the STRB for the typical crustal thickness outside of the STRB
of ~60 km [Neumann et al., 2008], which we assume represents the crustal thickness within the western STRB
prior to the inferred extension. For a typical Basin and Range strain rate of a 5 × 1015 s1, the total force change
is positive at the lower end of expected Martian thermal gradients (~15–20 K/km), indicating that delocalization
effects are greater than the localization effects and thus predicting wide rifting. At the higher end of expected
thermal gradients (20–35 K/km), the total force change transitions back to negative values, predicting core complex formation (Figure 12c). Strain rates of 1017–1014 s1 were tested, spanning the range from low values that
are likely applicable to Mars to higher values typical of wide rifting on Earth. For the range of estimated thermal
gradients for the Noachian period of 15 to 35 K/km [Schultz and Watters, 2001; McGovern et al., 2002, 2004; Grott
et al., 2005], wide rifting can occur for all strain rates >1017 s1. Increasing the strain rate shifts the wide to narrow rift transition to higher thermal gradients for a given crustal thickness (Figure 12d). For the expected thermal
gradients on Mars (>15 K/km) and typical Basin and Range strain rates (2 × 1016–5 × 1015 s1, [Sonder and
Jones, 1999]; wide rifting or core complex formation is still preferred (Figure 12d).
An approximate estimate of the strain rate in the western STRB can be made based on the constraints on the
timing and total strain. Since the formation of the western STRB is constrained to be before the NoachianHesperian transition at ~3.7 Ga, [Schultz and Tanaka, 1994], 800 Myr is an upper bound on the duration of
extension. Extension beginning coevally with the deformation of basins in western Terra Sirenum at
~4.0 Ga would suggest a duration of 300 Myr. If the observed extension began prior to or in the early stages
of Tharsis development during and subsequent to the Middle Noachian at ~3.9 Ga [Dohm et al., 2001c, 2007],
this would suggest a more stringent upper bound of ~200 Myr. If we take the total strain as 0.09 (from the
crustal thinning) or 0.18 (from the shortening of the western Terra Sirenum craters), the lower bounds on
the strain rate range from 3.6 × 1018 to 2.9 × 1017 s1 for a formation timescale of 200–800 Myr. An upper
bound for the strain rates can be obtained by taking the typical tectonic velocities on Earth (~1 cm/yr) and
dividing by the total width of the western STRB (1050 km), yielding a strain rate of 3.0 × 1016 s1. For the
estimated thermal gradients of Noachian Mars, these strain rates predict either wide rift or core complex
mode extension. Narrow rifting would require strain rates greater than 1014 s1 and/or thermal gradients
lower than 15 K/km, unless some additional localizing effect is acting on the system (see discussion below).
Thus, the formation of a Basin and Range-style wide rift on Mars is an expected outcome of regional extension of the lithosphere for a wide range of likely conditions on early Mars.
As our thermal modeling results predicted wide rift extension for typical Martian conditions, it is important
to discuss the conditions that can lead to the formation of narrow rifts, which are common on Mars [Grott
et al., 2005; Hauber and Kronberg, 2005; Grott et al., 2007; Hauber et al., 2010]. The thermal gradient, strain,
and crustal thickness are the most important factors determining the mode of rifting. In order for narrow
rifts to form, higher strain rates, lower thermal gradients or a thinner crust is required [Grott, 2005]. The
presence of narrow rifts within Tharsis suggests high strain rates, since both a thin crust and a low thermal
gradient conflict with the expectations for a volcanically thickened magmatic complex. Alternatively,
narrow rifting at Coracis Fossae has been explained by invoking a thinner crust at the time of extension,
requiring a later stage of crustal thickening after the formation of rifts [Grott, 2005]. Narrow Martian rifts
have also been explained by a combination of high strain rates in the initial stages of rifting, together with
localized mechanical weakening by magmatism along the rift [Grott et al., 2007; Hauber et al., 2010]. Our
findings suggest that typical low strain rates for the Noachian away from areas of magmatic weakening
should be characterized by wide rifting.
6. Geodynamic Implications of the STRB
6.1. Compressional Interpretation of the STRB
The formation of the STRB has important implications for the tectonic and geodynamic history of Mars, but
the nature of those implications depends on the interpretation of the ridge belt. A compressional origin
for the ridges, as proposed by Schultz and Tanaka [1994], would require an unknown source of large
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
935
Journal of Geophysical Research: Planets
10.1002/2015JE004936
compressional stress radial to Tharsis in order to produce the compression in the STRB. Tharsis loading
stresses alone should cause strong circumferential extension along with weaker radial extension of the
lithosphere in this region [Banerdt et al., 1992; Banerdt and Golombek, 2000]. The orientations of the ridges
are consistent with the orientations of the younger Hesperian wrinkle ridges in the region, which have been
explained by the addition of an isotropic horizontal compressional stress arising from global thermal cooling
and contraction, augmented by the advection of both heat and lava from the interior to the surface during
the formation of Tharsis [Watters, 1993; Golombek et al., 2001; Andrews-Hanna et al., 2008].
To form the STRB by compressional tectonism [Schultz and Tanaka, 1994], the contractional strain must be
sufficient to first overcome the Tharsis-induced extension in the region and then accumulate to the value
of ~0.003 arrived at earlier by treating the STRB ridges as being bordered by symmetric thrust faults.
However, models of the global thermal and volcanic evolution [Hauck and Phillips, 2002] and constraints from
populations of strike-slip faults [Andrews-Hanna et al., 2008] indicate a total accumulated contractional strain
of <0.002 by the end of the Noachian. Even lower amounts of contractional strain are predicted based on a
global inventory of faults [Nahm and Schultz, 2011]. Given the need to overcome the local extensional stresses from Tharsis loading, the expected contractional strain was insufficient to form the western STRB by compressional tectonism. Tharsis development at the end of the Noachian varied from the present-day
configuration based on geologic mapping investigations. By the Late Noachian, major development of the
Ceraunius rise, Syria Planum, the Thaumasia and Tempe plateaus, and Arsia rise (beneath the Tharsis
Montes) had occurred but only incipient to no development of Alba Mons, Olympus Mons, and Tharsis
Montes [Dohm et al., 2001c, 2007; Anderson et al., 2001; Isherwood et al., 2013]. However, the long-wavelength
shape of Tharsis was likely in place at this time. The progression from graben formation, to strike-slip faulting,
to wrinkle ridge formation in this region around the Noachian-Hesperian boundary [Andrews-Hanna et al.,
2008] indicates that stresses west of Tharsis were not yet compressional during the formation of the STRB
earlier in the Noachian.
A compressional origin for the STRB would require transitions from radial compression early in the Noachian
during the formation of the STRB, to circumferential extension at later stages to form the radial graben
swarms, to strike-slip faulting, and finally a return to circumferential compression to form the wrinkle ridges.
Such a scenario might have been brought about by a rapid period of volcanic outpouring during early Tharsis
construction causing an early period of rapid global contraction, with the gradual addition of membraneflexural stresses from the growing volcanic load causing a transition to strike-slip and extensional faulting
in the region. The later continued effects of global cooling and contraction could drive renewed thrust
faulting to form the wrinkle ridges. However, this scenario is not compatible with other evidence for the
magmatic-tectonic evolution of Tharsis [Anderson et al., 2001], and the magnitude of the compressional
strains required to form the western STRB by compression would be difficult to overcome with the addition
of membrane-flexural strains. Alternatively, an early phase of isostatic compensation of the rise could cause
radial compressional strains outside the rise [Banerdt et al., 1992]. However, a transition from Airy isostatic to
Tharsis’ present-day state of membrane-flexural compensation is not expected.
The distribution of the ridges in a narrow belt at the edge of Tharsis requires that either the strain itself or the
release of the strain would need to have been somehow localized within the STRB. For a compressional origin, the observed crustal thinning within the western STRB makes it necessary to invoke a thinner crust prior
to the formation of the ridges, with that thinning caused by an independent and unknown mechanism.
However, this preexisting thinner crust would result in a stronger lithosphere due to the stronger coupling
of the mantle and crustal lithosphere [Montési and Zuber, 2003b; Grott and Breuer, 2008], which could impede
the formation of the ridge belt by compression. Thus, a compressional origin for the western STRB is at odds
with the crustal thinning in the region and is difficult to reconcile with the strain history of Tharsis.
6.2. Extensional Interpretation of the STRB
An extensional formation mechanism for the STRB leads to different implications for early Martian geodynamics. Although the Basin and Range Province of the western United States is likely a by-product of plate
tectonics on the Earth, the immediate cause of the structures is a result of extensional strains and not
necessarily dependent on a plate tectonic environment. Both wide and narrow rifts can form on Mars from
local conditions independent of plate tectonics, and we explore such scenarios here. The orientations of
the ridges approximately concentric to Tharsis are not consistent with the orientation of extensional
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
936
Journal of Geophysical Research: Planets
10.1002/2015JE004936
membrane-flexural strains from Tharsis loading, which should cause radial normal faulting [Banerdt et al.,
1992; Banerdt and Golombek, 2000]. However, a pattern of circumferential normal faulting is consistent
with the expected radial extension arising from uplift of the lithosphere over a giant mantle plume or super
swell, for which membrane stresses dominate over flexural stresses [Banerdt et al., 1992]. This pattern is the
opposite of that expected for smaller regions of uplift, in which bending stresses dominate and radial normal faulting is expected. For Mars, membrane stresses exceed the flexural stresses for spherical harmonic
degrees <8 [Turcotte et al., 1981], corresponding to half wavelengths >1250 km. A best fit circle to the arc
of the STRB yields a diameter of 2730 km, confirming that it would be dominated by membrane stresses.
Evidence for an early phase of Tharsis uplift is found in the presence of ancient terrain and magnetic
anomalies high on Tharsis along the Claritas Fossae, interpreted as evidence of uplifted pre-Tharsis basement, as well as ancient Tharsis-concentric structures in Claritas Fossae [Phillips et al., 1990; Johnson and
Phillips, 2005]. The magnetic signature of the basement was likely destroyed by Tharsis volcanism farther
to the east and northeast [Lillis et al., 2009]. Geologic investigation by Dohm et al. [2009, 2015] has
resulted in the identification of a collection of stratigraphic, geomorphic, structural, geochemical, and
geophysical features which point to ancient dynamic activity including the formation of the Claritas
Rise, which could record pre-Tharsis and/or incipient Tharsis activity. Previous studies have suggested
that Tharsis may have formed in response to a giant spherical harmonic degree-one mantle plume
[Harder and Christensen, 1996], as has also been proposed for Earth to explain the two large-scale mantle
upwellings under the south Pacific and Africa [Maruyama, 1994]. The hypothesized superplume would
have resulted in the formation of igneous plateaus, volcanoes, fault systems, and lava flow fields [Dohm
et al., 2001c, 2007; Baker et al., 2007; Mège and Masson, 1996]. Individual upwellings have been interpreted
to have resulted in centers of magmatic and tectonic activity such as Claritas, Syria Planum, and Alba
Mons [Anderson et al., 2001].
Alternatively, other workers have argued for an origin through Tharsis as a result of slab rollback and lithospheric
subduction triggered by the Argyre impact [Yin, 2012b]. Early work suggested that Tharsis volcanism may have
been a passive effect of a preexisting thin and fractured lithosphere in this region that favored volcanic activity
[Solomon and Head, 1982]. However, both of the above hypotheses have difficulty explaining the thick lithosphere required to support Tharsis [Phillips et al., 2001]. Tharsis formation likely entailed a component of permanent uplift resulting from intrusive activity [Phillips et al., 1990], though a mantle upwelling is the most likely cause
of that intrusive activity. Once Tharsis had formed, continued volcanic activity could be passively driven as a result
of the insulating effect of the volcanically thickened crust [Schumacher and Breuer, 2007].
The early stages of Tharsis formation may thus have involved plume-induced uplift of the lithosphere,
before the rise transitioned into a membrane-flexurally supported load, leading to an evolving stress
regime in the area. It is possible that extension triggered by plume-induced uplift in the incipient stages
of Tharsis formation may have generated radial extensional stresses consistent with the orientations of
the STRB ridges. The circumferential faulting in the STRB, as opposed to the radial triple junction rifts attributed to plume-induced uplift on Earth and Venus [Mège and Ernst, 2001], would then be a result of the larger size of the hypothesized Martian plume relative to the size of the planet. Uplift of the lithosphere above
a flattened plume head [Campbell and Griffiths, 1990] would also tend to focus the extension at the margins
of the area of uplift.
STRB extension driven by a mantle plume would be consistent with hypotheses that a mantle plume is
responsible for extension across the Basin and Range of the western U.S. [Parsons et al., 1994; Parsons,
1995]. If the STRB formed as a result of membrane-flexural radial extension arising from plume-induced uplift,
the center of the plume would have been located at the center of a circular arc fit to the ridge belt. This would
place the plume center somewhat southwest of the present-day center of Tharsis (19°N, 236°E), qualitatively consistent with a proposed plume migration path from the south [Hynek et al., 2011]. The implied
migration of the plume would alter the stress regime, which may help to explain the en-echelon pattern of
some of the ridges. A later transition to volcanic loading would cause circumferential extension outside of
Tharsis, which is consistent with the radial graben observed in the area. Then, Hesperian-aged strike-slip
faults and circumferential wrinkle ridges would develop as a result of the superimposed stresses from global
cooling and contraction [Watters, 1993; Golombek et al., 2001; Andrews-Hanna et al., 2008]. The transition to
compression could also allow reactivation of normal faults in a thrusting sense, making it difficult to interpret
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
937
Journal of Geophysical Research: Planets
10.1002/2015JE004936
the tectonic style of the ridge belt. Reactivation of pre-Tharsis basement structures at any stage in the evolution
of the region would complicate the interpretation further.
The biggest challenge to the extensional origin is the generation of the inferred strain. The amount of extensional strain in the western STRB calculated from the crustal thinning of up to 0.09 is difficult to account for
on a single-plate planet such as Mars. Strains from membrane and flexural deformation are generally low, and
thus, simple plume-induced uplift cannot account for significant net extension of the lithosphere. Inversions
of the present-day global gravity and topography suggest extensional strains less than 0.005 in areas of
loading and flexure [Banerdt and Golombek, 2000]. The total strain that can be generated by simple flexural
uplift can be approximated by considering the resulting plate curvature. For an assumed uplift over a region
of known diameter, the radius of curvature can be calculated from the chord length of a circle. The strain is
then calculated as
y
εxx ¼ ;
(9)
R
where y is half the lithospheric thickness and R is the radius of curvature of the uplift. For uplift of 1–10 km
over a region 2730 km in diameter, the predicted strain ranges from 1 × 105 to 1 × 104 for a lithosphere
thickness of 20 km or 5 × 105 to 5 × 104 for a lithosphere thickness of 100 km. The flexural strain would
be increased by treating the uplift as a pancake shape, rather than a broad dome, using the width of western
STRB (1050 km) to calculate the radius of curvature. Uplift of 1–10 km then leads to strains ranging from
2 × 105 to 2 × 104 for a lithosphere thickness of 20 km or 9 × 105 to 9 × 104 for a lithosphere thickness
of 100 km. These flexural strains are >100 times lower than the strain of 0.09 inferred from the crustal
thickness, which requires an additional extensional mechanism in order to generate the strain in the area.
If an added source of extensional strain is invoked to explain the formation of the western STRB, it becomes
necessary to accommodate this extension with compression elsewhere, especially if Mars is considered to
have lacked plate tectonics throughout its evolution. As discussed in section 4.4, the craters of western
Terra Sirenum display significant ductile deformation, comparable in magnitude to the extension in the western STRB. However, the crust in the region of the deformed craters is not thickened as would be expected
from compression. In the next section, we propose a scenario linking the tectonic evolution of Tharsis, the
western STRB in northern Terra Sirenum, and western Terra Sirenum.
6.3. Proposed Tectonic and Geodynamic Evolution
We here propose a possible scenario to explain the inferred Basin and Range-style extension in the STRB and
the adjacent ductile shortening of craters in western Terra Sirenum. This scenario is consistent with both
observations of tectonism and ductile strain in these regions, as well as geodynamic models. Wide rift style
tectonic extension of the STRB may have been triggered by plume-induced uplift of the Tharsis region prior
to the construction of the present-day volcanic plateau, which would have driven limited radial extension
within the STRB. Simultaneously, downwelling mantle flow beneath western Terra Sirenum could have led
to a thickening of the crust through viscous coupling of the mantle with the ductile lower crust.
Parameterized convection models [Morschhauser et al., 2011] have shown that conditions expected for early
Mars could have led to lower crustal delamination and recycling, aided by the basalt-ecologite transition in
areas of thickened crust. Intense deformation of the overlying upper crust would be expected to result from
the delamination and recycling of the lower crust beneath Terra Sirenum. This is consistent with the observed
population of heavily deformed basins. Significant deformation observed in these basins appears to have a
major ductile component, which may have been encouraged by the thicker crust and higher heat flow prior
to lower crustal recycling. This upper crustal shortening and lower crustal recycling in western Terra Sirenum
would have enabled wide rift style extension to continue in the STRB. Later, radial extension within STRB
would be replaced by circumferential extension due to the transition to volcanic loading and membraneflexural support of Tharsis. The resulting membrane-flexural stresses would have led to the formation of
Tharsis-centered graben swarms such as the Memnonia and Sirenum Fossae that transect the western
STRB in our study region [Anderson et al., 2001], likely aided by dike propagation from Tharsis [Schultz
et al., 2004; Wilson and Head, 2002; Ernst et al., 2001]. Gradual cooling and contraction of Mars during the
Hesperian would have increased the horizontal compressional stresses leading to the formation of strike-slip
faults and circumferential thrust faults due to radial shortening within the western STRB [Andrews-Hanna
et al., 2008].
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
938
Journal of Geophysical Research: Planets
10.1002/2015JE004936
We emphasize that the above scenario is speculative. An extensional origin for the STRB cannot be confirmed
by our geological and geophysical analyses. The nature and cause of the significant ductile deformation in
western Terra Sirenum is unclear. Furthermore, the geodynamic evolution of Tharsis is uncertain, including
both the early stages of the formation of the rise and the activity that predates the hypothesized incipient
Tharsis upwelling [Schultz et al., 2004]. Nevertheless, many aspects of the tectonic and geodynamic evolution of
the STRB can be explained by the scenario proposed above. Regardless of the aforementioned ambiguities, it is
clear that the STRB and western Terra Sirenum together record a period of intense deformation of the crust early
in Martian history. Although the possible operation of plate tectonics on early Mars remains controversial, tectonic
evidence supports the possibility of significant horizontal strains, ductile deformation, and (more speculatively)
lower crustal recycling on early Mars. Unlike the more traditional fully stagnant lids of the Moon or Mercury, it
appears that at least some regions of Mars may have experienced a semimobile lithosphere.
7. Summary and Discussion
The South Tharsis Ridge Belt, which remains one of the most poorly understood ancient tectonic features of
Mars, is composed of 29 ridges, with typical width of 60 km, length of 250 km, and height of 1.5 km [Schultz
and Tanaka, 1994]. Schultz and Tanaka [1994] interpreted the ridges to be a belt related to compressional
tectonism and buckling of the lithosphere related to the development of Tharsis. A compressional origin
for the STRB is supported primarily by the morphology and the arcuate-like pattern of the ridges with respect
to Tharsis [Schultz and Tanaka, 1994], as well as by the observation of shortening of some craters in the
region. However, a compressional interpretation of the ridges would present challenges to our current
understanding of the stress history of Tharsis and Mars as a whole, requiring a strong source of either
isotropic or Tharsis-radial compressional stress early in Martian history.
Acknowledgments
We are grateful to An Yin, Amanda Nahm,
and an anonymous reviewer for their
thorough and thoughtful reviews. This
work was supported by grants to J.C.A.H.
from the NASA Mars Fundamental
Research Program (NNX10AM91G) and
Planetary Geology and Geophysics
Program (NNX14AO75G). J.M.D. was
supported by both the JSPS KAKENHI
grant 26106002 (Hadean BioScience
(Grant-in-Aid for Scientific Research on
Innovative Areas)) and the Tokyo Dome
Corporation regarding the TeNQ exhibit
and the branch of Space Exploration
Education & Discovery, the University
Museum, the University of Tokyo. The
MOLA data were obtained from NASA
Planetary Data System (http://pdsgeosciences.wustl.edu/missions/mgs),
THEMIS data were obtained from
(http://www.mars.asu.edu/data/), and
National elevation data set data were
obtained from (http://ned.usgs.gov/).
The data used in section 4.4 are available in the supporting information.
KARASOZEN ET AL.
Alternatively, we note that the morphology and morphometry of both the individual ridges and the STRB as a
whole bear a strong resemblance to the Basin and Range extensional tectonic province of the western United
States. The interpretation of the STRB as an extensional tectonic province is supported by the associated crustal thinning, as well as by possible examples of elongated craters. The wide spacing of the ridges in the STRB
in comparison to the ridges in the Basin and Range and other examples of distributed extension can be
explained as a consequence of the thicker lithosphere on Mars. Although most Martian rifts are of the narrow
rift type, typical conditions on early Mars would have favored wide rifting as inferred to have occurred at the
STRB. This Tharsis-radial extension may have been initiated by an early period of plume-induced uplift during
the incipient stages of Tharsis formation, though membrane-flexural deformation of the lithosphere above a
mantle plume alone cannot produce the extensional strain inferred from the crustal thinning. Significant
extension across the STRB would need to have been accommodated by compression elsewhere. We speculate that the heavily deformed craters of Western Terra Sirenum may have undergone compression in parallel
with the extension in the STRB, possibly driven by convergent lower crustal flow and delamination.
Although the available evidence cannot conclusively rule in favor of either extensional or compressional
models of formation, our analysis favors the extensional interpretation. If correct, this would make the
STRB the only example of documented Basin and Range-style extension on Mars. Regardless of the interpretation, the STRB clearly represents a mode of tectonism not observed elsewhere on Mars. The extensive
tectonic deformation within the STRB, and the possibly related deformation that pervades the adjacent
western Terra Sirenum province, provides evidence of large strains and a partially mobile lithosphere early
in Martian history.
References
Aittola, M., T. Öhman, J. J. Leitner, and J. Raitala (2007), The characteristics of polygonal impact craters on Venus, Earth Moon Planets,
101(1–2), 41–53.
Anderson, F. S., and R. R. Grimm (1998), Rift processes at the Valles Marineris, Mars: Constraints from gravity on necking and rate-dependent
strength evolution, J. Geophys. Res., 103, 11,113–11,124, doi:10.1029/98JE00740.
Anderson, R. C., J. M. Dohm, M. P. Golombek, F. C. Albert, B. J. Franklin, K. L. Tanaka, J. Lias, B. Peer, and H. Early (2001), Primary centers and
secondary concentrations of tectonic activity through time in the western hemisphere of Mars, J. Geophys. Res., 106, 20,563–20,585,
doi:10.1029/2000JE001278.
Anderson, R. C., J. M. Dohm, S. Robbins, B. Hynek, and J. Andrews-Hanna (2012), Terra Sirenum: Window into Pre-Tharsis and Tharsis phases
of Mars evolution, Lunar Planet. Sci., 43, Abstract 2803.
Andrews-Hanna, J. C. (2012a), The formation of Valles Marineris: 1. Tectonic architecture and the relative roles of extension and subsidence,
J. Geophys. Res., 117, E03006, doi:10.1029/2011JE003953.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
939
Journal of Geophysical Research: Planets
10.1002/2015JE004936
Andrews-Hanna, J. C. (2012b), The formation of Valles Marineris: 2. Stress focusing along the buried dichotomy boundary, J. Geophys. Res.,
117, E04009, doi:10.1029/2011JE003954.
Andrews-Hanna, J. C. (2012c), The formation of Valles Marineris: 3. Super-isostasy, stress, sedimentation, and subsidence, J. Geophys. Res.,
117, E06002, doi:10.1029/2012JE004059.
Andrews-Hanna, J. C., and M. T. Zuber (2010), Elliptical craters and basins on the terrestrial planets, Geol. Soc. Am., 465, 1–13, doi:10.1130/
2010.2465(01).
Andrews-Hanna, J. C., M. T. Zuber, and S. A. Hauck II (2008), Strike-slip faults on Mars: Observations and implications for global tectonics and
geodynamics, J. Geophys. Res., 113, E08002, doi:10.1029/2007JE002980.
Anguita, F., A. F. Farelo, V. López, C. Mas, M. J. Muñoz-Espadas, A. Márquez, and J. Ruiz (2001), Tharsis dome, Mars: New evidence for
Noachian-Hesperian thick-skin and Amazonian thin-skin Tectonics, J. Geophys. Res., 106, 7577–7589, doi:10.1029/2000JE001246.
Atwater, T. (1970), Implications of plate tectonics for the Cenozoic tectonic evolution of western North America, Geol. Soc. Am. Bull., 81(12),
3513–3536, doi:10.1130/0016-7606(1970)81[3513:IOPTFT]2.0.CO;2.
Atwater, T., and J. Stock (1998), Pacific-North America plate tectonics of the Neogene southwestern United States—An update, Int. Geol. Rev.,
40(5), 375–402, doi:10.1080/00206819809465216.
Axen, G. J., W. J. Taylor, and J. M. Bartley (1993), Space-time patterns and tectonic controls of Tertiary extension and magmatism in the Great
Basin of the western United States, Geol. Soc. Am. Bull., 105(1), 56–76, doi:10.1130/0016-7606(1993)105<0056:STPATC>2.3.CO;2.
Baker, V. R., S. Maruyama, and J. M. Dohm (2007), Tharsis superplume and the geological evolution of early Mars, in Super-Plumes: Beyond
Plate Tectonics, edited by D. A. Yuen et al., pp. 507–523, Springer, Netherlands.
Banerdt, W. B., and M. P. Golombek (2000), Tectonics of the Tharsis region of Mars: Insights from MGS topography and gravity, Lunar Planet. Sci.,
31, 2038.
Banerdt, W. B., M. P. Golombek, and K. L. Tanaka (1992), Stress and tectonics on Mars, in Mars, edited by H. H. Kieffer et al., pp. 249–297,
Univ. of Arizona Press, Tucson, Ariz.
Bennett, R., J. Davis, and B. Wernicke (1999), Present-day pattern of Cordilleran deformation in the Western United States, Geology, 27, 371–374.
Bennett, R., B. Wernicke, N. Niemi, A. Friedrich, and J. Davis (2003), Contemporary strain rates in the northern Basin and Range province from
GPS data, Tectonics, 22(2), 1008, doi:10.1029/2001TC001355.
Bennett, R., S. Hreinsdóttir, M. S. Velasco, and N. Fay (2007), GPS constraints on vertical crustal motion in the northern Basin and Range,
Geophys. Res. Lett., 34, L22319, doi:10.1029/2007GL031515.
Bland, M. T., and A. P. Showman (2007), The formation of Ganymede’s grooved terrain: Numerical modeling of extensional necking instabilities,
Icarus, 189, 439–456, doi:10.1016/j.icarus.2007.01.012.
Bland, M. T., W. B. McKinnon, and A. P. Showman (2010), The effects of strain localization on the formation of Ganymede’s grooved terrain,
Icarus, 210, 396–410, doi:10.1016/j.icarus.2010.06.008.
Bottke, W. F., S. G. Love, D. Tytell, and T. Glotch (2000), Interpreting the elliptical crater populations on Mars, Venus, and the Moon, Icarus, 145,
108–121, doi:10.1006/icar.1999.6323.
Buck, W. R. (1988), Flexural rotation of normal fault, Tectonics, 7, 959–973, doi:10.1029/TC007i005p00959.
Buck, W. R. (1991), Modes of continental lithospheric extension, J. Geophys. Res., 96, 20,161–20,178, doi:10.1029/91JB01485.
Buck, W. R. (2007), Dynamic processes in extensional and compressional settings: The dynamics of continental breakup and extension,
in Crust and Lithosphere Dynamics: Treatise on Geophysics, Crustal and Lithosphere Dynamics, vol. 6, edited by A. B. Watts, pp. 335–376,
Elsevier, Amsterdam.
Campbell, I. H., and R. W. Griffiths (1990), Implications of mantle plume structure for the evolution of flood basalts, Earth Planet. Sci. Lett.,
99, 79–93.
Caristan, Y. (1982), The transition from high temperature creep to fracture in Maryland diabase, J. Geophys. Res., 87, 6781–6790, doi:10.1029/
JB087iB08p06781.
Choi, E. S., and M. Gurnis (2003), Deformation in transcurrent and extensional environments with widely spaced weak zones, Geophys. Res.
Lett., 30(2), 1076, doi:10.1029/2002GL016129.
Coney, P. J., and T. A. Harms (1984), Cordilleran metamorphic core complexes: Cenozoic extensional relics of Mesozoic compression, Geology,
12(9), 550–54, doi:10.1130/0091-7613(1984)12<550:CMCCCE>2.0.CO;2.
Connerney, J. E. P., M. H. Acuña, N. F. Ness, G. Kletetschka, D. L. Mitchell, R. P. Lin, and H. Reme (2005), Tectonic implications of Mars crustal
magnetism, Proc. Natl. Acad. Sci. U.S.A., 102, 14,970–14,975.
Constenius, K. (1996), Late Paleogene extensional collapse of the Cordilleran foreland fold and thrust belt, Geol. Soc. Am. Bull., 108, 20–39.
Dickinson, W. R. (2002), The Basin and Range province as a composite extensional domain, Intl. Geol. Rev., 44, 1–38, doi:10.2747/0020-6814.44.1.1.
Dixon, T. H., M. Miller, F. Farina, H. Wang, and D. Johnson (2000), Present-day motion of the Sierra Nevada block and some tectonic implications
for the Basin and Range province, North America Cordillera, Tectonics, 19, 1–24, doi:10.1029/1998TC001088.
Dohm, J. M., and K. L. Tanaka (1999), Geology of the Thaumasia region, Mars: Plateau development, valley origins, and magmatic evolution,
Planet. Space Sci., 47, 411–431.
Dohm, J. M., S. Maruyama, and V. R. Baker (2001a), Mars plate tectonics: Surface geology and analyses of topographic and geophysical data,
Abstract P32C-0561 presented at 2011 Fall Meeting, AGU, San Francisco, Calif., Dec.
Dohm, J. M., K. L. Tanaka, and T. M. Hare (2001b), Geologic map of the Thaumasia region of Mars, U.S. Geol. Surv. Misc. Invest. Ser. Map, I-2650.
Dohm, J. M., J. C. Ferris, V. R. Baker, R. C. Anderson, T. M. Hare, R. G. Strom, N. G. Barlow, K. L. Tanaka, J. E. Klemaszewski, and D. H. Scott (2001c),
Ancient drainage basin of the Tharsis region, Mars: Potential source for outflow channel systems and putative oceans or paleolakes,
J. Geophys. Res., 106, 32,943–32,958, doi:10.1029/2000JE001468.
Dohm, J. M., S. Maruyama, V. R. Baker, R. C. Anderson, J. C. Ferris, and T. M. Hare (2002), Plate tectonism on early Mars: Diverse geological and
geophysical evidence, Lunar Planet. Sci. Conf. XXXIII, Abstract 1639.
Dohm, J. M., S. Maruyama, V. R. Baker, and R. C. Anderson (2007), Traits and evolution of the Tharsis superplume, Mars, in Superplumes: Beyond
plate Tectonics, edited by D. A. Yuen et al., pp. 523–537, Springer, Netherlands.
Dohm, J. M., et al. (2009), Claritas rise, Mars: Pre-Tharsis magmatism?, J. Volcanol. Geotherm. Res., 185, 139–156.
Dohm, J. M., et al. (2013), Mars evolution, in Mars: Evolution, Geology, and Exploration, edited by A. G. Fairén, pp. 1–34, Nova Sci. Inc., Hauppauge,
New York.
Dohm, J. M., et al. (2015), The Mars plate-tectonic-basement hypothesis, Lunar Planet. Sci., Conf. XXXXVI, Abstract 1741.
Dombard, A. J., and W. B. McKinnon (2001), Formation of grooved terrain on Ganymede: Extensional instability mediated by cold, superplastic
creep, Icarus, 154, 321–336, doi:10.1006/icar.2001.6728.
Doser, D. I., and R. B. Smith (1989), An assessment of source parameters of earthquakes in the Cordillera of the western United States,
Bull. Seismol. Soc. Am., 79, 1383–1409.
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
940
Journal of Geophysical Research: Planets
10.1002/2015JE004936
Ellis, M., A. Densmore, and R. Anderson (1999), Development of mountainous topography in the Basin Ranges USA, Basin Res., 11, 21–42,
doi:10.1046/j.1365-2117.1999.00087.x.
Ernst, R. E., E. B. Grosfils, and D. Mège (2001), Giant dike swarms: Earth, Venus, and Mars, Annu. Rev. Earth Planet. Sci., 29, 489–534, doi:10.1146/
annurev.earth.29.1.489.
Fairén, A. G., and J. M. Dohm (2004), Age and origin of the lowlands of Mars, Icarus, 168, 277–284, doi:10.1016/j.Icarus.2003.11.025.
Fairén, A. G., J. Ruiz, and F. Anguita (2002), An origin for the linear magnetic anomalies on Mars through accretion of terranes: Implications for
dynamo timing, Icarus, 160, 220–223, doi:10.1006/icar.2002.6942.
Fielder, G. (1965), Lunar Geology, Lutterworth Press, London.
Fink, J. H., and R. C. Fletcher (1981), Variations in thickness of Ganymede’s lithosphere determined by spacings of lineations, Lunar Planet. Sci. Conf.,
12, 277–278.
Fleitout, L., and C. Froidevaux (1983), Tectonic stresses in the lithosphere, Tectonics, 2(3), 315–324, doi:10.1029/TC002i003p00315.
Fletcher, R. C., and B. Hallet (1983), Unstable extension of the lithosphere: A mechanical model for basin-and-range structure, J. Geophys. Res.,
88, 7457–7466, doi:10.1029/JB088iB09p07457.
Floran, R. J., and M. R. Dence (1976), Morphology of the Manicouagan ring-structure, Quebec, and some comparisons with lunar basins and
craters, Lunar Planet. Sci. Conf., 7, 2845–2865.
Friedrich, A., B. Wernicke, and N. Niemi (2003), Comparison of geodetic and geologic data from the Wasatch region, Utah, and implications for
the spectral character of Earth deformation at periods of 10 to 10 million years, J. Geophys. Res., 108(B4), 2199, doi:10.1029/2001JB000682.
Gesch, D. B. (2007), The national elevation dataset, in Digital Elevation Model Technologies and Applications: The DEM User’s Manual, 2nd ed.,
edited by D. Maune, pp. 99–118, Am. Soc. for Photogramm. and Remote Sens., Bethesda, Md.
Gesch, D., M. Oimoen, S. Greenlee, C. Nelson, M. Steuck, and D. Tyler (2002), The national elevation dataset, Photogramm. Eng. Remote Sens.,
68, 5–32.
Ghent, R. (2002), Ribbon spacing in Venusian tessera: Implications for layer thickness and thermal state, Geophys. Res. Lett., 29(20), 2000,
doi:10.1029/2002GL015994.
Ghent, R., and V. Hansen (1999), Structural and kinematic analysis of eastern Ovda Regio,Venus: Implications for crustal plateau formation,
Icarus, 139, 116–136, doi:10.1006/icar.1999.6085.
Glazner, A. F., and J. M. Bartley (1984), Timing and tectonic setting of Tertiary low-angle normal faulting and associated magmatism in the
southwestern United States, Tectonics, 3(3), 385–396, doi:10.1029/TC003i003p00385.
Golombek, M. P., F. S. Anderson, and M. T. Zuber (2001), Martian wrinkle ridge topography: Evidence for subsurface faults from MOLA,
J. Geophys. Res., 106, 23,811–23,821, doi:10.1029/2000JE001308.
Golombek, M. P., et al. (2006), Erosion rates at the Mars Exploration Rover landing sites and long-term climate change on Mars, J. Geophys. Res.,
111, E12S10, doi:10.1029/2006JE002754.
Gomes, R., H. F. Levison, K. Tsiganis, and A. Morbidelli (2005), Origin of the cataclysmic Late Heavy Bombardment period of the terrestrial
planets, Nature, 435, 466–469, doi:10.1038/nature03676.
Grimm, R. E., and S. W. Squyres (1985), Spectral analysis of groove spacing on Ganymede, J. Geophys. Res., 90, 2013–2021, doi:10.1029/
JB090iB02p02013.
Grott, M. (2005), Late crustal growth on Mars: Evidence from lithospheric extension, Geophys. Res. Lett., 32, L23201, doi:10.1029/2005GL024492.
Grott, M., and D. Breuer (2008), The evolution of the Martian elastic lithosphere and implications for crustal and mantle rheology, Icarus, 193,
503–515, doi:10.1016/j.Icarus.2007.08.015.
Grott, M., and D. Breuer (2010), On the spatial variability of the Martian elastic lithosphere thickness: Evidence for mantle plumes?, J. Geophys. Res.,
115, E03005, doi:10.1029/2009JE003456.
Grott, M., E. Hauber, S. C. Werner, P. Kronberg, and G. Neukum (2005), High heat flux on ancient Mars: Evidence from rift flank uplift at Coracis
Fossae, Geophys. Res. Lett., 32, L2120, doi:10.1029/2005GL023894.
Grott, M., P. Kronberg, E. Hauber, and B. Cailleau (2007), Formation of the double rift system in the Thaumasia Highlands, Mars, J. Geophys. Res.,
112, E06006, doi:10.1029/2006JE002800.
Guest, J. E., and J. B. Murray (1969), Nature and origin of Tsiolkovsky crater, lunar farside, Planet. Space Sci., 17(1), 121–141.
Hamblin, W. K. (1965), Origin of ‘reverse drag’ on the downthrown side of normal faults, Geol. Soc. Am. Bull., 76, 1145–1164.
Hamilton, W. B. (2005), Palinpastic reconstruction in Nevada by reversal extensional deformation, Rep. prepared for Newmont Mining Company.
Hamilton, W. G., and W. B. Myers (1966), Cenozoic tectonics of the western United States, Rev. Geophys., 4, 509–549, doi:10.1029/
RG004i004p00509.
Hansen, V. (1998), Ribbon terrain formation, southwestern Fortuna Tessera, Venus: Implications for lithosphere evolution, Icarus, 132,
321–343, doi:10.1006/icar.1998.5897.
Harder, H., and U. R. Christensen (1996), A one-plume model of Martian mantle convection, Nature, 380, 507–509, doi:10.1038/380507a0.
Hartmann, W., and G. Neukum (2001), Cratering chronology and the evolution of Mars, Space Sci. Rev., 96, 165–194, doi:10.1023/A:1011945222010.
Hashimoto, G. L., M. Roos-Serote, S. Sugita, M. S. Gilmore, L. W. Kamp, R. W. Carlson, and K. H. Baines (2009), Felsic highland crust on Venus
suggested by Galileo Near-Infrared Mapping Spectrometer data, J. Geophys. Res., 114, 1–10, doi:10.1029/2008JE003134.
Hauber, E., and P. Kronberg (2005), The large Thaumasia graben on Mars: Is it a rift?, J. Geophys. Res., 110, E07003, doi:10.1029/2005JE002407.
Hauber, E., M. Grott, and P. Kronberg (2010), Martian rifts: Structural geology and geophysics, Earth Planet. Sci. Lett., 294, 393–410,
doi:10.1016/j.epsl.2009.11.005.
Hauck, S. A., and R. J. Phillips (2002), Thermal and crustal evolution of Mars, J. Geophys. Res., 107(E7), 5052, doi:10.1029/2001JE001801.
Hoogenboom, T., and S. E. Smrekar (2006), Elastic thickness estimates for the northern lowlands of Mars, Earth Planet. Sci. Lett., 248, 830–839,
doi:10.1016/j.epsl.2006.06.035.
Hynek, B. M., S. J. Robbins, O. Šrámek, and S. J. Zhong (2011), Geological evidence for a migrating Tharsis plume on early Mars, Earth Planet.
Sci. Lett., 310, 327–333.
Isherwood, R. J., L. M. Jozwiak, J. C. Jansen, and J. C. Andrews-Hanna (2013), The volcanic history of Olympus Mons from paleo-topography
and flexural modeling, Earth Planet. Sci. Lett., 363, 88–96.
Jackson, J. A., and N. J. White (1989), Normal faulting in the upper continental crust: Observations from regions of active extension, J. Struct.
Geol., 11, 15–36, doi:10.1016/0191-8141(89)90033-3.
Johnson, C. L., and R. J. Phillips (2005), Evolution of the Tharsis region of Mars: Insights from magnetic field observation, Earth Planet. Sci. Lett.,
230, 241–254.
Jones, C. H., J. R. Unruh, and L. J. Sonder (1996), The role of gravitational potential energy in active deformation in the southwestern United
States, Nature, 381, 37–41, doi:10.1038/381037a0.
Karato, S., and P. Wu (1993), Rheology of the upper mantle: A synthesis, Science, 260, 771–778.
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
941
Journal of Geophysical Research: Planets
10.1002/2015JE004936
Karato, S. I., M. S. Paterson, and J. D. Fitzgerald (1986), Rheology of synthetic olivine aggregates—Influence of grain size and water, J. Geophys. Res.,
91, 8151–8176, doi:10.1029/JB091iB08p08151.
Kiefer, W. S., and M. S. Johnson (1995), Physical properties controlling the style of extensional deformation in the Tharsis region of Mars,
Lunar Planet. Sci., 26, 741–742.
Kogan, L., S. Fisseha, and R. Bendick (2012), Lithospheric strength and strain localization in continental extension from observations of the
East African Rift, J. Geophys. Res., 117(3), 1–16, doi:10.1029/2011JB008516.
Kohlstedt, D. L., B. Evans, and S. J. Mackwell (1995), Strength of the lithosphere: Constraints imposed by laboratory experiments, J. Geophys. Res.,
100(B9), 17,587–17,602, doi:10.1029/95JB01460.
Kopal, J. (1966), An Introduction to the Study of the Moon, D. Reidel Co., Norwell, Mass.
Lillis, R. J., J. Dufek, J. E. Bleacher, and M. Magna (2009), Demagnetization of crust by magmatic intrusion near the Arsia Mons volcano:
Magnetic and thermal implications for the development of the Tharsis province, J. Volcanol. Geotherm. Res., 185, 123–138.
Liu, L., and D. R. Stegman (2011), Segmentation of the Farallon slab, Earth Planet. Sci. Lett., 311, 1–10, doi:10.1016/j.epsl.2011.09.027.
Mackwell, S. J., M. E. Zimmerman, and D. L. Kohlstedt (1998), High-temperature deformation of dry diabase with application to tectonics on Venus,
J. Geophys. Res., 103, 975–984, doi:10.1029/97JB02671.
Maruyama, S. (1994), Plume tectonics, J. Geol. Soc. Jpn., 100, 24–49.
McGovern, P. J., S. C. Solomon, D. E. Smith, M. T. Zuber, M. Simons, M. A. Wieczorek, R. J. Phillips, G. A. Neumann, O. Aharonson, and J. W. Head (2002),
Localized gravity/topography admittance and correlation spectra on Mars: Implications for regional and global evolution, J. Geophys. Res.,
107(E12), 5136, doi:10.1029/2002JE001854.
McGovern, P. J., S. C. Solomon, D. E. Smith, M. T. Zuber, M. Simons, M. A. Wieczorek, R. J. Phillips, G. A. Neumann, O. Aharonson, and J. W. Head
(2004), Correction to “Localized gravity/topography admittances and correlation spectra on Mars: Implications for regional and global
evolution”, J. Geophys. Res., 109, E07007, doi:10.1029/2004JE002286.
McQuarrie, N., and B. P. Wernicke (2005), An animated tectonic reconstruction of southwestern North America since 36 Ma, Geosphere, 1(3),
147–172, doi:10.1130/GES00016.1.
Mège, D., and P. Masson (1996), A plume tectonics model for the Tharsis province, Mars, Planet. Space Sci., 44, 1499–1546, doi:10.1016/
S0032-0633(96)00113-4.
Mège, D., and R. E. Ernst (2001), Contractional effects of mantle plumes on Earth, Mars and Venus, in Mantle Plumes: Their Identification Through
Time, Geol. Soc. Am. Spec. Pap., vol. 352, edited by R. E. Ernst and K. L. Buchan, pp. 103–140.
Mohapatra, G. K., and R. A. Johnson (1998), Localization of listric faults at thrust ramps beneath the Great Salt Lake Basin, Utah: Evidence from
seismic imaging and finite element modeling, J. Geophys. Res., 103(B5), 10,047–10,063, doi:10.1029/98JB00023.
Montési, L. G. J., and M. T. Zuber (2003a), Spacing of faults at the scale of the lithosphere and localization instability: 2. Application to the
Central Indian Basin, J. Geophys. Res., 108(B2), 2111, doi:10.1029/2002JB001924.
Montési, L. G. J., and M. T. Zuber (2003b), Clues to the lithospheric structure of Mars from wrinkle ridge sets and localization instability,
J. Geophys. Res., 108(E6), 5048, doi:10.1029/2002JE001974.
Montési, L. G. J., and M. T. Zuber (2003c), Spacing of faults at the scale of the lithosphere and localization instability: 1. Theory, J. Geophys. Res.,
108(B2), 2110, doi:10.1029/2002JB001923.
Montgomery, D. R., S. M. Som, M. P. A. Jackson, B. C. Schreiber, A. R. Gillespie, and J. B. Adams (2009), Continental-scale salt tectonics on Mars
and the origin of Valles Marineris and associated outflow channels, Geol. Soc. Am. Bull., 121, 117–133.
Moore, J. G. (1960), Curvature of normal faults in the Basin and Range province of the western United States, U.S. Geol. Surv. Prof. Pap. 400–B,
409–411.
Morschhauser, A., M. Grott, and D. Breuer (2011), Crustal recycling, mantle dehydration, and the thermal evolution of Mars, Icarus, 212(2),
541–558, doi:10.1016/j.Icarus.2010.12.028.
Mueller, N., J. Helbert, G. L. Hashimoto, C. C. C. Tsang, S. Erard, G. Piccioni, and P. Drossart (2009), Venus surface thermal emission at 1 μm
in VIRTIS imaging observations: Evidence for variation of crust and mantle differentiation conditions, J. Geophys. Res., 114, E00B17,
doi:10.1029/2008JE003118.
Mueller, S., and R. J. Phillips (1995), On the reliability of lithospheric constraints derived from models of outer-rise flexure, Geophys. J. Int., 123, 887–902.
Nahm, A. L., and R. A. Schultz (2010), Evaluation of the orogenic belt hypothesis for the formation of the Thaumasia Highlands, Mars,
J. Geophys. Res., 115, E04008, doi:10.1029/2009JE003327.
Nahm, A. L., and R. A. Schultz (2011), Magnitude of global contraction on Mars from analysis of surface faults: Implications for Martian thermal
history, Icarus, 211, 389–400, doi:10.1016/j.icarus.2010.11.003.
Neumann, G. A., F. G. Lemonie, D. E. Smith, and M. T. Zuber (2008), Megacrust3—Crustal thickness inversion from recent MRO gravity solutions,
Lunar Planet. Sci., XXXIX, Abstract 2167.
Niemi, N., B. Wernicke, A. Frederick, M. Simmons, R. Bennett, and J. Davis (2004), BARGEN continuous GPS data across the eartern Basin and
Range province, and implications for fault system, Geophys. J. Int., 159, 842–862, doi:10.1111/j.1365-246X.2004.02454.x.
Nimmo, F. (2004), Dynamics of rifting and modes of extension on icy satellites, J. Geophys. Res., 109, E01003, doi:10.1029/2003JE002168.
Öhman, T., M. Aittola, V. P. Kostama, J. Raitala, and J. Korteniemi (2008), Polygonal impact craters in Argyre region, Mars: Implications for
geology and cratering mechanics, Meteorit. Planet. Sci., 43, 1605–1628.
Pappalardo, R. T., and G. C. Collins (2005), Strained craters on Ganymede, J. Struct. Geol., 27(5), 827–838, doi:10.1016/j.jsg.2004.11.010.
Parsons, T. (1995), The Basin and Range Province, in Continental Rifts: Evolution, Structure and Tectonics, edited by K. Olsen, pp. 277–324,
Elsevier, Amsterdam.
Parsons, T., G. A. Thompson, and N. H. Sleep (1994), Mantle plume influence on Neogene uplift and extension of the U.S. western Cordillera?,
Geology, 22, 83–86.
Phillips, R. J., N. M. Sleep, and W. B. Banerdt (1990), Permanent uplift in magmatic systems with application to the Tharsis region of Mars,
J. Geophys. Res., 95, 5089–5100, doi:10.1029/JB095iB04p05089.
Phillips, R. J., et al. (2001), Ancient geodynamics and global-scale hydrology on Mars, Science, 291, 2587–2591.
Quaide, W. L., D. E. Gault, and R. A. Schmidt (1965), Gravitative effects on lunar impact structures, Ann. N. Y. Acad. Sci., 123, 641–655.
Ruiz, J. (2014), The early heat loss evolution of Mars and their implications for internal and environmental history, Sci. Rep., 4, 4338,
doi:10.1038/srep04338.
Saylor, J., P. DeCelles, G. Gehrels, M. Murphy, R. Zhang, and P. Kapp (2010), Basin formation in the High Himalaya by arc-parallel extension and
tectonic damming: Zhada basin, southwestern Tibet, Tectonics, 29, TC1004, doi:10.1029/2008TC002390.
Scholz, C. H. (2002), Earthquakes and Faulting, 2nd ed., Cambridge Univ. Press, Cambridge, U. K.
Schultz, R. A. (2000), Localization of bedding plane slip and backthrust faults above blind thrust faults: Keys to wrinkle ridge structure,
J. Geophys. Res., 105(E5), 12,035–12,052, doi:10.1029/1999JE001212.
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
942
Journal of Geophysical Research: Planets
10.1002/2015JE004936
Schultz, R. A., and K. L. Tanaka (1994), Lithospheric-scale buckling and thrust structures on Mars: The Coprates rise and south Tharsis ridge
belt, J. Geophys. Res., 99(E4), 8371–8385, doi:10.1029/94JE00277.
Schultz, R. A., and T. R. Watters (2001), Forward mechanical modeling of the Amenthes Rupes thrust fault on Mars, Geophys. Res. Lett., 28(24),
4659–4662, doi:10.1029/2001GL013468.
Schultz, R. A., C. H. Okubo, C. L. Goudy, and S. J. Wilkins (2004), Igneous dikes on Mars revealed by Mars Orbiter Laser Altimeter topography,
Geology, 32, 889–892.
Schumacher, S., and D. Breuer (2007), An alternative mechanism for recent volcanism on Mars, Geophys. Res. Lett., 34, L14202,
doi:10.1029/2007GL030083.
Severinghaus, J., and T. Atwater (1990), Cenozoic geometry and thermal state of the subducting slabs beneath western North America, in Basin
and Range Extension Tectonics Near the Latitude of Las Vegas, Nevada, vol. 176, edited by B. P. Wernicke, pp. 1–22, Geol. Soc. Am. Mem.,
doi:10.1130/MEM176-p1.
Shoemaker, E. M. (1963), Impact mechanics at Meteor crater, Arizona, in The Moon, Meteorites, and Comets, edited by B. M. Middlehurst and
G. P. Kuiper, pp. 301–336, Univ. of Chicago Press, Chicago.
Sleep, N. H. (1994), Martian plate tectonics, J. Geophys. Res., 99, 5639–5655, doi:10.1029/94JE00216.
Smith, D. E., et al. (2001), Mars Orbiter Laser Altimeter: Experiment summary after the first year of global mapping of Mars, J. Geophys. Res.,
106, 23,689–23,722, doi:10.1029/2000JE001364.
Solomon, S. C., and J. W. Head (1982), Evolution of the Tharsis province of Mars—The importance of heterogeneous lithospheric thickness
and volcanic construction, J. Geophys. Res., 87, 9755–9774, doi:10.1029/JB087iB12p09755.
Sonder, L. J., and C. H. Jones (1999), Western United States extension: How the West was widened, Annu. Rev. Earth Planet. Sci., 27, 417–462,
doi:10.1146/annurev.earth.27.1.417.
Sonder, L. J., P. C. England, B. P. Wernicke, and R. L. Christianse (1987), A physical model for Cenozoic extension of western North America,
in Continental Extensional Tectonics, edited by M. P. Coward, J. F. Dewey, and P. L. Hancock, Geol. Soc. Spec. Publ., 28, 187–201,
doi:10.1144/GSL.SP.1987.028.01.14.
Spudis, P. D. (1993), The Geology of Multi-Ring Impact Basins, Camb. Planet. Sci. Ser., vol. 8, 263 pp., Cambridge Univ. Press, New York.
Stewart, J. H. (1980), Regional tilt patterns of late Cenozoic basin-range fault blocks, western United States, Geol. Soc. Am. Bull., 91, 460–464.
Tanaka, K. L., and M. P. Golombek (1994), Strain measurements of impact craters on Tempe Terra, Mars, Lunar Planet. Sci., XXV, 1377-1378.
Tanaka, K. L., J. A. Skinner, J. M. Dohm, R. P. Rossman III, E. J. Kolb, C. M. Fortezzo, T. Platz, G. G. Michael, and T. M. Hare (2014), Geologic map of
Mars, U.S. Geol. Surv. Sci. Invest. Map, 3292.
Turcotte, D. L., R. J. Willemann, W. F. Haxby, and J. Norberry (1981), Role of membrane stresses in the support of planetary topography,
J. Geophys. Res., 86(B5), 3951–3959, doi:10.1029/JB086iB05p03951.
Velasco, M. S., R. A. Bennett, R. A. Johnson, and S. Hreinsdóttir (2010), Subsurface fault geometries and crustal extension in the eastern Basin
and Range Province, western U.S., Tectonophysics, 488, 131–142, doi:10.1016/j.tecto.2009.05.010.
Watters, T. R. (1993), Compressional tectonism on Mars, J. Geophys. Res., 98(E9), 17,049–17,060, doi:10.1029/93JE01138.
Watters, T. R. (2003), Thrust faults along the dichotomy boundary in the eastern hemisphere of Mars, J. Geophys. Res., 108(E6), 1–12,
doi:10.1029/2002JE001934.
Watters, T. R., and M. S. Robinson (1997), Radar and photoclinometric studies of wrinkle ridges on Mars, J. Geophys. Res., 102, 10,889–10,903,
doi:10.1029/97JE00411.
Werner, S. C. (2008), The early Martian evolution—Constraints from basin formation ages, Icarus, 195(1), 45–60, doi:10.1016/j.icarus.2007.12.008.
Wernicke, B. (1992), Cenozoic extensional tectonics of the U.S. Cordillera, in The Cordilleran Orogen: Conterminous U.S. The Geology of North
America, G-3, edited by B. C. Burchfiel, P. W. Lipman, and M. L. Zoback, pp. 553–581, Geol. Soc. Am., Boulder, Colo.
Wernicke, B., and G. J. Axen (1988), On the role of isostasy in the evolution of normal fault systems, Geology, 16(9), 848–851.
Wernicke, B., and B. C. Burchfiel (1982), Modes of extensional tectonics, J. Struct. Geol., 4, 105–115.
Wernicke, B. P., R. L. Christiansen, P. C. England, and L. J. Sonder (1987), Tectonomagmatic evolution of Cenozoic extension in the North American
Cordillera, in Continental Extensional Tectonics, Geol. Soc. Spec. Publ., vol. 28, edited by M. P. Coward, J. F. Dewey, and P. L. Hancock,
pp. 203–221.
Wernicke, B., G. J. Axen, and J. K. Snow (1988), Basin and Range extensional tectonics at the latitude of Las Vegas, Nevada, Geol. Soc. Am. Bull.,
100(11), 1738–1757.
Wieczorek, M. A., and M. T. Zuber (2004), Thickness of the Martian crust: Improved constraints from geoid-to-topography ratios, J. Geophys. Res.,
109, E01009, doi:10.1029/2003JE002153.
Wilkins, S. J., R. A. Schultz, R. C. Anderson, J. M. Dohm, and N. H. Dawers (2002), Deformation rates from faulting at the Tempe Terra extensional
province, Mars, Geophys. Res. Lett., 29(18), 1884, doi:10.1029/2002GL015391.
Wilson, L., and J. W. Head (2002), Tharsis-radial graben systems as the surface manifestation of plume-related dike intrusion complexes:
Models and implications, J. Geophys. Res., 107(E08), 5057, doi:10.1029/2001JE001593.
Yin, A. (2012a), Structural analysis of the Valles Marineris fault zone: Possible evidence for large-scale strike-slip faulting on Mars, Lithosphere,
4(4), 286–330, doi:10.1130/L192.1.
Yin, A. (2012b), An episodic slab-rollback model for the origin of the Tharsis rise on Mars: Implications for initiation of local plate subduction
and final unification of a kinematically linked global plate-tectonic network on Earth, Lithosphere, 4(6), 553–593, doi:10.1130/L195.1.
Zoback, M. L., R. E. Anderson, and G. A. Thompson (1981), Cainozoic evolution of the state of stress and style of tectonism of the Basin and
Range Province of the western United States, Philos. Trans. R. Soc., A, 300(1454), 407–434, doi:10.1098/rsta.1981.0073.
Zuber, M. T., E. M. Parmentier, and R. C. Fletcher (1986), Extension of continental lithosphere: A model for two scales of Basin and Range
deformation, J. Geophys. Res., 91(B5), 4826–4838, doi:10.1029/JB091iB05p04826.
Zuber, M. T., et al. (2000), Internal structure and early thermal evolution of Mars from Mars Global Surveyor topography and gravity, Science,
287(5459), 1788–1793.
Zuber, M. T., F. G. Lemoine, D. E. Smith, A. S. Konopliv, S. E. Smrekar, and S. W. Asmar (2007), Mars Reconnaissance Orbiter Radio Science
Gravity Investigation, J. Geophys. Res., 112, E05S07, doi:10.1029/2006JE002833.
KARASOZEN ET AL.
FORMATION OF THE SOUTH THARSIS RIDGE BELT
943