THE FORMATION MECHANISM of SOUTH THARSIS RIDGE BELT, MARS by Ezgi Karasözen

Home , List of terrae on Mars, Lithospheric mantle, Mars, Rupes

THE FORMATION MECHANISM OF SOUTH THARSIS RIDGE BELT, MARS

Ezgi Karasözen A thesis submitted to the Faculty and the Board of Trustees of the Colorado School of Mines in partial fulfillment of the requirements for the degree of Master of Science (Geophysics).

Golden, Colorado

Date ______

Signed: ______

Ezgi Karasözen

Signed: ______

Dr. Jeffrey C. Andrews-Hanna

Thesis Advisor

Golden, Colorado

Date ______

Signed:______

Dr. Terence K. Young Professor and Head Department of Geophysics

ABSTRACT

The South Tharsis Ridge Belt (STRB) partially surrounds Tharsis in an arc around the southwest part of the rise. It consists of 29 large ridges separated by distances up to 110 km, with average relief of 1.5 km above the surrounding plains. Because the STRB is among the oldest tectonic features associated with Tharsis, it may provide key information on the early evolution of Tharsis and the ancient tectonic history of Mars. Earlier studies concluded that the ridges formed through compressional tectonism by a combination of a buckling instability and thrust faulting. However, when the shape, size, and separation of the ridges are considered, the STRB resembles the extensional Basin and Range province on Earth, both of which are characterized by series of parallel mountain ranges separated by broad valleys. In this thesis, we evaluate both extensional and compressional hypotheses for the origin of the ridges using evidence from topographic profiles, deformed craters, boundary element modeling, crustal thickness models, and thermal modeling. Though no single model explains all aspects of the ridges, the topography of the ridges and crustal thinning associated with the belt support an extensional origin. Thermal modeling results show that Basin and Range-style wide rifting in the STRB under expected thermal gradients of 15-20 K/km should occur for strain rates comparable to those in the Basin and Range. We suggest that the large amount of extensional strain inferred in the STRB may have been accommodated by the significant ductile deformation observed in the western Terra

Sirenum craters. This has significant implications for the incipient development of Tharsis and possibly pre-Tharsis phases of Mars’ evolution. However, with the available evidence we cannot unequivocally reject a compressional mechanism for forming the STRB or confirm an extensional mechanism.

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TABLE OF CONTENTS

ABSTRACT ...... iii

LIST OF FIGURES ...... vi

LIST OF TABLES ...... ix

ACKNOWLEDGMENTS ...... x

CHAPTER 1 INTRODUCTION ...... 1

CHAPTER 2 ANALOGS ...... 5

2.1 Compressional Analog: Amenthes Rupes ...... 5

2.2. Extensional Analog: Basin and Range ...... 6

2.2.1 Tectonics of basin and range extension ...... 9

CHAPTER 3 TOPOGRAPHIC ANALYIS AND TECTONIC MODELING ...... 12

3.1 Topographic Analysis ...... 12

3.1.1 Results ...... 13

3.2 Tectonic Modeling ...... 16

3.2.1 Results ...... 17

CHAPTER 4 STRAIN INDICATORS ...... 22

4.1 Deformed craters of the South Tharsis Ridge Belt ...... 22

4.2 Strain from crustal thickness ...... 25

4.3 Strain from horizontal tectonic deformation ...... 26

4.4 Strain from simple flexure...... 28

4.5 Strain from craters of Western Terra Sirenum ...... 31

4.5.1 Strain from Rose Diagrams ...... 31

4.5.2 Timing of the deformation ...... 35

CHAPTER 5 THERMAL MODELING ...... 37

5.1 Theory and Method ...... 39

5.2 Crustal buoyancy force...... 41

5.3 Yield strength ...... 42

5.4 Results ...... 43

CHAPTER 6 DISCUSSION AND CONCLUSIONS ...... 47

6.1 Compressional origin ...... 47

6.2 Extensional origin ...... 48

6.3 Proposed tectonic and geodynamic evolution ...... 50

REFERENCES CITED ...... 52

APPENDIX A SLOPE CALCULATION ...... 63

APPENDIX B THERMAL MODELING ...... 73

LIST OF FIGURES

Figure 1.1 MOLA topography map (Smith et al., 2001) showing the South Tharsis Ridge Belt (Figure 1.2a) and Western Terra Sirenum (Figure 4.5), in cylindrical projection. The altitudes above 4.69 km are not discriminated. Please note the varying labeled altitude intervals which are due to the selected classification method (quantile classification) in ARCMAP...... 2

Figure 1.2 Topographic maps of (a,b) the South Tharsis Ridge belt on Mars, (c) the Basin and Range province on Earth and (d) Amentes Rupes on Mars. The lines correspond to the locations of the profiles in Figure 3.1. The letters with boxes in (a) correspond to the locations of the craters in Figure 4.1. All maps are in a cylindrical projection, and panel c uses the NAD83 datum. Latitude - longitude range of each panel is: (a) 10-39°S, 210-240°E (b) 12-19°S, 212- 218°E (c) 38.0-41.5°N, 113-118°W (d) 0-4°N, 109-113°E...... 4

Figure 3.1. Representative profiles of ridges in the (a, b, c) South Tharsis Ridge Belt, (d, e) Basin and Range, and (f) Amenthes Rupes. Locations of the profiles are shown in Figure 1.2. The vertical exaggeration is 40 ××× in (a-f). Profiles g, h, and i reproduce the profiles in (a,b,c), (d,e) and f at 1x vertical exaggeration and identical horizontal scales...... 15

Table 3.1 Topographic measurements of the South Tharsis Ridge Belt (STRB), Basin and Range province (BR), and Amenthes Rupes (AR). These values are calculated from the profiles in Appendix A...... 16

Figure 3.2 Cross sections (a,b) and topographic profiles of (c, e) symmetric and (d, f) secondary back thrust modeling, respectively. Vertical exaggeration in (c) is 160x, (d) is 200x, and (d, e) is 1x...... 19

Figure 3.3 The crossing depth for symmetric thrust faults bordering the ridges as a function of slope for different ridge widths...... 21

Figure 4.1 Impact craters in the STRB, in cylindrical projection. Capital letters denote the crater names. See Figure 1.2a for the locations of craters...... 24

Figure 4.2 (a) Crustal thickness map of STRB based on the model of (Neumann et al., 2008). (b) Representative crustal thickness profile, (c) topographic map and (d) representative topographic profile of STRB. Dashed lines indicate the typical crustal thickness and elevation within STRB. Location of the representative profile is shown in (a) and (c) with white line. Cylindrical projection is used in the topographic map...... 27

Figure 4.3 Strain versus dip angle for (a) STRB and (b) Basin and Range...... 28

Figure 4.4 Craters of Western Terra Sirenum, marked craters are the 7 largest basins with diameters >170km, in Mercator projection. Latitude – longitude range is 15 -55°S – 175-215°E ...... 32

Figure 4.5 Largest basins (a) location of the crater with crater rim points (black dots) and best-fit circle (white dashed lines) in Mercator projection, (b) Rose diagrams showing extensional (red) and compressional (blue) strain values as a function of azimuth relative to the best-fit circle right). See Figure 4.4 for the crater numbers in panel (a)...... 33

Figure 4.6 Average strain Rose diagrams (a) large 7 basins (b) all craters diameters < 170 km (10 degree intervals)...... 35

Figure 4.7 Crater size frequency distributions of all crater population...... 36

Table 5.1 Model Parameters of Mars...... 40

Table 5.2 Rheological parameters assumed for Mars. A is a material dependent constant and is the activation energy and n is a material dependent exponent. (1) (Buck, 1991), (2) (Grott 2005; Grott and Breuer 2008)...... 43

Figure 5.1 Rifting mode diagrams for strain rate 8x10 -15 s-1 (a) 1x10 -16 s-1 (b). Total forces for a 60 km thick crust and the strain rate 5 ×××10 -15 s-1s-1 (c) and semi-log rifting mode diagram (d) for a 60 km thick crust. The range of thermal gradients estimated for the Noachian period is indicated by the shading in (c) and the range of strain rates estimated for the Basin and Range (2 ×××10 -16 – 5×××10 -15 s-1) is indicated by the shading in (d)...... 45

Figure A-1. Locations of the individual profiles of the South Tharsis Ridge Belt, in cylindrical projection. All of the profiles are taken from West to East. Red profiles show the locations of representative profiles for this region. Latitude – longitude range is -12,-32 - 205,222...... 63

Figure A-2. Individual profiles of the South Tharsis Ridge Belt plotted all together (left) and separately (right). The vertical exaggeration is 20××× in all panels. See figure A-1 for the locations of the profiles. Please note the variable scales within the profiles...... 64

Figure A-3. Locations of the individual profiles of the Basin and Range, and datum of the map is NAD83. All of the profiles are taken from West to East. Red profiles show the locations of representative profiles for this region. Latitude - longitude range is 38,41.5N - 113,118W...... 66

Figure A-4. Individual profiles of the Basin and Range plotted separately (left) and all together (right). The vertical exaggeration is 20 ××× in all panels. See figure A-3 for the locations of the profiles. Please note the variable scales within the profiles...... 67

Figure A-5. Locations of the individual profiles of the Amenthes Rupes, in cylindrical projection. All of the profiles are taken from West to East. Red profiles show the location of representative profile for this region. Latitude – longitude range is 0,4 - 108,112...... 72

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Figure A-6. Individual profiles of the Amenthes Rupes plotted all together (left) and separately (right). The vertical exaggeration is 20 ××× in all panels. See figure A-5 for the locations of the profiles. Please note the variable scales within the profiles...... 72

Figure B-1. Brittle-ductile depth change for different rheologies...... 74

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LIST OF TABLES

Table 3.1 Topographic measurements of the South Tharsis Ridge Belt (STRB), Basin and Range province (BR), and Amenthes Rupes (AR). These values are calculated from the profiles in Appendix A...... 16

Table 5.1 Model Parameters of Mars...... 40

ACKNOWLEDGMENTS

I would like to express my gratitude to my advisor Dr. Jeffrey Andrews-Hanna for his guidance, critical vision, motivation and kindness through all of the stages of this master thesis. I also want to thank my committee members, Dr. Thomas Davis, and Dr. Warren Hamilton, for their advice, support and valuable input to my research.

Special thanks are extended to Dr. Kadri Dagdelen for his invaluable guidance and support throughout my graduate school experience in the Colorado School of Mines.

I would like to thank all of the Planetary group members, current and former, especially

Brian Davis for his friendship, valuable comments and helpful discussions.

I would also like to thank to my “Golden” family; Basak Sener Kaya, Tugce Conger, Basar

Basbug, Hakan Corapcioglu and Murat Yuksel Kaya for being my family here. It would be impossible to finish this research without them. Also, I would like to acknowledge all members of TSA for their friendship and continuous motivation.

I would like to thank to my extended family, my dear F.R.I.E.N.D.S, who managed to be with me in all stages of my research, from overseas; Ezgi Tincer, Ayse Merve Genc, Cigdem

Cankara and Gizem Tincer.

I would like to thank to Onur Conger, for always finding a way to make me smile and making my life joyful, colorful and livable.

Last but not the least important, I am deeply grateful to my dearest parents Zerrin and

Bulent Karasozen, for everything. I would not be able to succeed without them. They truly merit the dedication of this thesis.

To my parents Bulent & Zerrin

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CHAPTER 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, there is little geological or geophysical supporting evidence that plate tectonics played an active role in the history of Mars. Most of the tectonic features from Late Noachian onwards are consistent with the interaction of global contraction and loading/flexural stresses. They are characterized by small strains and faulting in a continuous lithosphere. However, a subset of ancient tectonic features of pre-Noachian to Early Noachian age cannot be explained by these processes, including rift zones (Hauber et al., 2010) and the Thaumasia highlands (Fairéna and

Dohm, 2004). These ancient features display significant strains that need to be accommodated somehow, given the one-plate nature of the martian lithosphere. Understanding the mechanism responsible for these strains may shed light on the nature of early Martian geodynamics.

One of these ancient tectonic features is the South Tharsis Ridge Belt (STRB) (Figure

1.1), which forms an arcuate shape around Tharsis (Schultz and Tanaka, 1994). The STRB is among the oldest tectonic features associated with Tharsis, but the mechanism of its formation is unclear. In this thesis, I discuss several models to understand the formation mechanism of the

STRB, which may provide key information on Tharsis’ early evolution and the ancient tectonic history of Mars.

The South Tharsis Ridge Belt consists of 29 large ridges (Schultz and Tanaka, 1994) separated by distances up to 110 km, with average relief of 1.5 km above the surrounding plains.

This ridge belt was first studied by Schultz and Tanaka (1994), who carried out a detailed

1 photogeologic mapping. They concluded that the ridges formed by a combination of a buckling instability and thrust faulting. However, the geomorphology of the ridges resembles the parallel mountain ranges of the Basin and Range province of the western United States. The Basin and

Range ridges 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).

In this thesis, I evaluate both extensional and compressional hypotheses for the origin of the ridges using evidence from topographic profiles, deformed craters, boundary element modeling, crustal thickness models and thermal modeling. I compare the western portion of the

STRB (Figure 1.2a, 1.2b) with the Basin and Range province of Earth (Figure 1.2c) as an extensional analog, and the Amenthes Rupes thrust fault on Mars (Figure 1.2d) as a compressional analog. In section 2, I discuss these analogs and their tectonic environments. In section 3, I compare topographic profiles of the STRB to the Basin and Range province and to

Amenthes Rupes, and use tectonic modeling to understand the fault geometry of the STRB. In section 4, the strain across the province is evaluated from craters of the STRB and from crustal thickness models, and then possible strain mechanisms are compared. In section 5, thermal modeling is carried out to understand the conditions of Noachian Mars that can result in the formation of the STRB through wide mode rifting. In section 6, possible formation scenarios of 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. This has significant implications for the incipient development of Tharsis and possibly pre-Tharsis phases of Mars’ evolution.

CHAPTER 2

ANALOGS

First, I discuss possible compressional and extensional analogs to the STRB, in order to compare it to better-characterized tectonic systems.

2.1 Compressional Analog: Amenthes Rupes

Schultz and Tanaka (1994) interpreted the STRB as a set of compressional ridges formed as a result of a buckling instability. During horizontal compression or extension, the lithosphere can become unstable, resulting in buckling or necking. A buckling instability predicts broad folds in which deformation is distributed. If the lithosphere weakens with deformation, it develops a localized instability resulting in regularly spaced faults (Montesi and Zuber, 2003b).

Montesi and Zuber (2003b) analyzed how buckling/necking and fault spacing can be related to each other by studying the behavior of lithosphere under different deformation conditions. They found that in compressional environments, the lithosphere can produce regularly spaced faults through either a buckling instability or a localization instability (Montesi and Zuber, 2003b).

The suggested analog for these two processes is the Central Indian Basin (Montesi and

Zuber, 2003c). The shortening in the Central Indian Basin created folds in the basement with a wavelength of 200 km and an amplitude of 2 km as a result of a buckling instability. The fault spacing of 7-11 km is controlled by localization instability where faults penetrate through to the upper mantle (Montesi and Zuber, 2003c). Although the Central Indian Basin is a clear example

5 of the effects of buckling and localization instabilities, its ocean floor topography bears no resemblance to the STRB and I do not consider it further as an analog.

For a possible compressional analog for the STRB, I chose to analyze the well-defined martian lobate scarp, Amenthes Rupes. Lobate scarps are compressional structures interpreted as surface-breaking thrust faults (Watters and Robinson, 1996). They do not typically occur in populations with regular spacing between the ridges on Mars. However, it has been suggested that surface-breaking thrust faults combined with a buckling instability may generate a population of uniformly spaced compressional ridges (Schultz and Tanaka, 1994). Amenthes

Rupes (Figure 1d) is located south of the dichotomy boundary to the east of the Isidis basin. The lobate scarp population that includes Amenthes Rupes offsets heavily-cratered terrain of late

Noachian age (Schultz and Watters, 2001), and itself formed in the Late Noachian to Early

Hesperian (Watters and Robinson, 1999). Amenthes Rupes is a 380 km long thrust fault scarp with a maximum relief of ~1.2 km (Watters, 2003). It reflects the characteristic morphologic elements of lobate scarps, with a relatively steeply sloping front scarp and a gently sloping back scarp (Watters, 2003). The depth of faulting is estimated to range between 25 and 30 km, and the dip angle is estimated to vary between 25° and 30° with an average displacement of 1.5 km

(Schultz and Watters, 2001). The horizontal shortening ranges between 0.24 to 2.6 km for fault dips of 25°, and the strain rate was estimated as 10 -17 s-1 or less (Watters and Robinson, 1999).

2.2. Extensional Analog: Basin and Range

I suggest that the geomorphology of the ridges of the STRB resembles the extensional

Basin and Range province in the western United States, which I discuss as an extensional analog.

The Basin and Range is a unique tectonic province characterized by a pattern of alternating basins and ranges. This region is formed as a result of Cenozoic extension in western North

America (Sonder et al., 1987). Its widest point is more than 900 km across from east to west, and it extends from Canada, through the western United States, and across much of Mexico. The

Basin and Range differs markedly from most narrow continental rifts, and is commonly described as a “wide rift”. The timing, amount, and spatial variation of the extension are well known (Sonder and Jones, 1999; Dickinson, 2002; Dickinson, 2006; Hamilton, 2005). The driving forces for extension are thought to be a combination of boundary forces due to plate tectonics, traction and normal forces applied to the base of the lithosphere, as well as buoyancy forces due to the horizontal gradients in density. The latter two are mostly responsible for driving extension in the Basin and Range, while plate boundary-driven extension is responsible for shear deformation in the western parts and past extension in the southern parts (Sonder and Jones,

1999). The Basin and Range is subdivided into the Northern (NBR), Central (CBR), and

Southern (SBR) sub-provinces (Sonder and Jones, 1999). The NBR has the highest average elevation, highest heat flow, and thinnest crust. The SBR has the lowest average elevation of the three sub-provinces, the lowest heat flow, and is tectonically the least active. The CBR has the greatest local relief and represents a transition from the NBR to the SBR, with large north-south gradients in regional topography, heat flow, and gravity. It is tectonically active, particularly in its western half.

Extension in the Basin and Range did not occur in the form of a uniform stretching. The direction and amount of the extension and strain vary with time and location. For example, the

Great Basin province of the Basin and Range (the subsection of the Northern and Central Basin and Range province which is located between the Wasatch and the Sierra Nevada Mountains)

7 doubled in width during the Cenozoic period. The total crustal extension estimates vary between

10-300% (Parsons, 1995; Hamilton and Myers, 1966; Zoback et al., 1981; Wernicke, 1992;

Sonder and Jones, 1999). It should be noted here that strain is not uniformly distributed, with some regions displaying extreme extension (eg. CBR 250%; Sonder and Jones, 1999) and some displaying minor extension (e.g., <10%; Parsons, 1995). Variation of strain with depth is also highly likely because the crust has the same thickness in extended and unextended terrains

(Sonder and Jones, 1999). Strain rates, like strains, are strongly variable both spatially and temporally, and are estimated to range from 0.2–5×10 -15 s-1 (Sonder and Jones, 1999).

The crustal deformation in the Basin and Range has been studied using various approaches. Fletcher and Hallet (1983) treated the lithosphere as a strong plastic layer overlying a viscous layer of the same density, and suggested that the regular spacing is primarily determined by the thickness of the strong layer. Zuber et. al. (1986) examined implications of a stratified lithosphere on large-scale extensional deformation, and found that the regular spacing of tilts and range domains of Basin and Range are controlled by necking instabilities that arise from strength stratification of the lithosphere. In the Basin and Range, buoyancy forces resulting from the thickened crust are the main source of the extensional strain (Sonder et al., 1987). The style of extension has been attributed to the balance between localizing and de-localizing effects, with wide rifting occurring when the de-localizing effects become larger than the localizing effects (Buck, 2007). This is discussed further in Section 5 in the discussion of the tectonic and geodynamic implications of possible wide rifting at the South Tharsis Ridge Belt.

2.2.1 Tectonics of basin and range extension

The topography in the Basin and Range varies at different scales due to the pervasive extension. The ranges are spaced at an average of 25-35 km apart with basins 10-20 km wide

(Parsons, 1995). Individual ranges can display either symmetric or asymmetric topography. This basin and range structure is most prominent in the Great Basin, but also appears in other regions of the Basin and Range. Extensional structures show a variety of fault geometries. Extensional faults have been grouped by Wernicke and Burchfiel (1982) into non-rotational and rotational faults. Rotational faults have planar fault geometries and develop as a result of extension and progressive rotation of geological features (beds). In non-rotational extension, extension takes place on planar or listric faults without rotation of geological features.

The Basin and Range is a result of a complicated geometry of extensional faulting, as has been discussed by several authors. According to Hamilton and Myers (1966), many ranges in the Basin and Range are dominated by block faulting and each block in the province is bounded one or both sides by zones of dip-slip normal faults. The bounding faults mostly dip at 40°-80° in the basin-ward direction. Moore (1960) and Hamblin (1965) suggested that faults of the Basin and Range should flatten at depth because of the general concavity in plan of most ranges.

Hamilton and Myers (1966) concluded that if these faults are listric, then this would limit the depth of normal faulting to less than 10 km.

Zoback at al. (1981) grouped the generalized models of Basin and Range structure into horst and graben, tilted block, and listric fault models. Stewart (1980) examined regional tilt patterns for major range blocks within the Basin and Range, and identified broad regions of consistent tilt. This regional tilt domains are compatible with the tilted block and listric fault

9 models but not with the horst and graben model. Zoback et al. (1981) concluded that the Basin and Range province probably includes many types of faults, including steep faults that become subhorizontal at depths of 4-5 km, and other faults that decrease in dip and merge with a subhorizontal zone of decoupling at 15 km depth.

Wernicke and Burchfiel (1982) disagree with the commonly accepted high-angle non- rotational faulting model, suggesting instead that the large amount of extension is accompanied by widespread imbricate normal-fault blocks and subjacent very low-angle normal faults. They do accept that listric faulting is common in Basin and Range, but argue that these faults are not the only contributor to the extension.

Earthquake focal mechanisms show that much of present-day seismicity of the Basin and

Range occurs on high angle faults breaking the surface with dips of 45°-60° (e.g. Doser and

Smith 1989; Jackson and White, 1989; Ellis et. al., 1999). The depth of faulting is mostly limited to 15 km, just above the brittle-ductile transition. According to Hamilton (2005), the widely accepted listric faulting model of the Basin and Range is incompatible with the seismic evidence.

Following the Buck (1988) model of fault rotation, Hamilton (2005) suggests that the listric faults in the Basin and Range are rotated, high-angle normal faults. Buck (1988) states that down-rotation of hanging walls and back-rotation of footwalls are necessary products of upper- crustal flexure due to the isostatic response of topographic loading caused by the fault offsets.

This flexure rotates the fault, decreases its dip upward and stops the slip on the old fault. This process causes a new fault of steeper dip to break upward on the hanging wall block.

Landscape evolution models of the Basin and Range carried out by Ellis et al. (1999) suggest that the interaction between faulting and surface processes has produced a number of distinctive morphologies in the Basin and Range province. Some of these ranges show clear

10 asymmetry whereas the others are more symmetric. Their modeling includes cases both with displacement on a single normal fault and with a symmetrical horst. They included this second symmetrical case from Dutton’s (1880) paper where he suggests that the central and northern

Basin and Range are accommodated by normal faults with opposing dips. Ellis et al. (1999) concluded that symmetrical initial conditions coupled with pre-existing topography and surface processes can reproduce the topography in the Basin and Range. The symmetric and asymmetric ranges of Basin and Range province makes it a good extensional analogue for the STRB, as will be examined further in the topographic comparison in the next section.

CHAPTER 3

TOPOGRAPHIC ANALYIS AND TECTONIC MODELING

Topographic analysis was carried out to out the similarities and differences between

STRB and its analogues. For further evaluation, tectonic modeling was carried out to understand the geometry of faulting in the STRB which is responsible for the topographic profiles. In this section the details and results of topographic analysis and tectonic modeling are given.

3.1 Topographic Analysis

I now use topographic profiles to provide a quantitative comparison of the South Tharsis

Ridge Belt to the compressional and extensional analogs. The topographic analysis was performed using Mars Orbiter Laser Altimeter (MOLA) topography data for Mars (Smith et al.,

2001), together with the National Elevation Dataset (NED) for Earth (Gesch, 2007; Gesch et. al.,

2002). I analyzed the western portion of the STRB (Figure 1.2 a,b), which contains six well- defined ridges, with typical lengths of 250 km, widths of 60 km, heights of 1.5 km and separations between ridges of up to 110 km. For comparison, I examined topographic profiles of

Basin and Range ridges and Amenthes Rupes lobate scarp as extensional and compressional analogs (Figure 1.2c, d).

Multiple topographic profiles were taken perpendicular to each ridge axis, approximately equally spaced but avoiding areas that are highly deformed or cratered. 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 were quantified.

3.1.1 Results

The STRB ridges are characterized by a range of morphologies. Some ridges exhibit symmetric slopes on either side of the ridges, with an average value of 0.14 (Table 3.1; Figure

3.1a, Profile A-A’). However, other ridges show marked asymmetries in slope (Figure 3.1b,

Profile B-B’), with an average slope value of the steeper flanks of 0.19 and gentler flanks of

0.12. The ridges commonly exhibit an elevation difference from one side of the ridge to the other, likely a result of the accumulation of lava on the Tharsis-facing side. 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 may be bound by symmetric faults on either side, similar to the horst structures inferred for some Basin and Range ridges. The asymmetric profiles of other STRB ridges are consistent with the ridges being bound by a single fault, as observed in the majority of the Basin and Range ridges. Some ridges within the STRB show crested tops (Figure 3.1c, Profile C-C’) as in the case of most of the Basin and Range ridges.

Topographic profiles of the Basin and Range ridges show crested tops, which appear to be the result of fluvial erosion of the ranges. Their slopes are mostly symmetric with values of

0.143 and -0.149, for west- and east-facing limbs, respectively (see Table 3.1; Figure 3.1d,

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. However, some ridges have markedly asymmetric profiles (Figure 3.1e, Profile Y-

Y’). The average slope value of steeper flanks is 0.15 and gentler flanks is 0.10 which is similar to STRB slope values. This states that both regions are composed of a combination of symmetric and asymmetric ridges.

In contrast, topographic profiles of Amenthes Rupes (Figure 3.1f, Profile Z-Z’) differ significantly from STRB and Basin and Range. They clearly reflect the basic morphologic elements of lobate scarps with relatively steep sloping scarp faces and gentle back slopes

(Watters et. al., 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.

From a purely topographic standpoint, the combination of symmetric and asymmetric ridges in the 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 on Mars (Golombek et al.,

2006). The scarp dips of the STRB and the Basin and Range are similar, while the STRB differs from Amenthes Rupes in the much lower dip of the back slope in the latter. Among three tectonic provinces, the STRB has the highest average ridge height and width of 1.5 and 60 km, respectively (Table 3.1). Basin and Range and Amenthes Rupes have similar average heights, whereas the width of Amenthes Rupes is twice the average width of Basin and Range, reflecting the lower dip angle of the back slope. The most significant difference between the STRB and the

Basin and Range is the ridge spacing (110 km and 25-35 km (Parsons, 1995), respectively).

This topographic analysis suggests that the STRB ridges are mostly similar to the Basin and Range ridges. The symmetric ridges common in the STRB may indicate that symmetric horst structures are more common there than in the Basin and Range, though it is difficult to infer tectonic geometry from topography alone. In contrast, Amenthes Rupes displays a strong topographic asymmetry, and therefore is not a good match to the STRB ridges. The aspect ratios

14 of the STRB and Basin and Range ridges are similar, while Amenthes Rupes is wider relative to its height.

Figure 3.1. Representative profiles of ridges in the (a, b, c) South Tharsis Ridge Belt, (d, e) Basin and Range, and (f) Amenthes Rupes. Locations of the profiles are shown in Figure 1.2. The vertical exaggeration is 40 × in (a-f). Profiles g, h, and i reproduce the profiles in (a,b,c), (d,e) and f at 1x vertical exaggeration and identical horizontal scales.

Table 3.1 Topographic measurements of the South Tharsis Ridge Belt (STRB), Basin and Range province (BR), and Amenthes Rupes (AR). These values are calculated from the profiles in Appendix A.

Average Slope Average Slope Average Average Ridge # of ridges West East Height (km) Width (km)

STRB 0.13±0.04 -0.14±0.03 1.5±0.4 59.0±15.9 5

BR 0.13±0.02 -0.13±0.02 0.9±0.3 34.8±7.9 13

AR 0.11 -0.01 1.1 71.7 1

3.2 Tectonic Modeling

Topographic profiles revealed the similarities and differences between STRB and its analogues. For further evaluation, I need to understand the fault geometry responsible for these topographic profiles. Therefore, tectonic modeling was carried out to match the surface topography with the related fault geometries, in order to better understand the geometry of faulting in the STRB.

The surface topography surrounding a fault is the combined result of the displacement across the fault and the elastic deformation of the surrounding medium (Cohen, 1999). I use the boundary element dislocation program COULOMB (Toda et. al., 2005; Lin and Stein 2004) to test the possibility of a compressional origin for the STRB ridges. However, Basin and Range- style tectonism requires significant ductile deformation of the lower crust, and thus cannot be evaluated using the elastic boundary element model. The near-surface topography resulting from slip on a fault was determined using this elastic dislocation modeling program, which has been

16 previously used to model a variety of terrestrial and planetary faults, such as lobate scarps

(Schultz and Watters, 2001) and wrinkle ridges (Watters, 2004) on Mars.

I assume a Poisson’s ratio of 0.25 and Young’s modulus of 80 GPa. The fault surface was defined as a rectangular plane of specified dip angle and vertical extent in the subsurface. The amount of displacement and sense of slip were specified along the fault, and resulting stresses and material displacements were determined by using the stress functions for an elastic halfspace

(Okada, 1992). A tapered elliptical slip distribution was used in the down-dip direction. The model parameters were iteratively adjusted to achieve a good fit between the predicted and observed topography.

3.2.1 Results

I first performed tectonic modeling of Amenthes Rupes, following the parameters used in

Schultz and Watters (2001). The resultant tectonic profile matched well and confirmed that lobate scarps are well-explained by a simple thrust-faulting origin. I then attempted to fit the

STRB topography by compressional tectonism. Single thrust faults were not found to be capable of reproducing the observed symmetric steep scarps. Therefore, I tested the possibility of both symmetric thrust faults, as well as paired primary and secondary back thrust faults as proposed by Okubo and Schultz (2004) for the formation of wrinkle ridges. In the latter mechanism, the primary thrust fault is the main driver of the formation of wrinkle ridges. This upward- propagating thrust fault slips in mechanically weak horizons of the crust, leading to the localization of secondary brittle strain and the formation of secondary back thrusts.

For the case of symmetric faults (Figure 3.2a) I assumed a fault dip angle of 30°, and down-dip tapering slip amount of 1.5 km. Here elliptical slip distribution is used to taper the slip in the down dip direction to zero at the lower fault tip. In order for faults not to cross-cut each other at depth, the depth of the faulting must be less than 17 km for a ridge width of 60 km. The resultant topographic profile (Figure 3.2c) is extremely different from the topographic profiles of

STRB. Varying the modeling parameters did not improve the fit. More realistic scenarios where the lower fault tips do not approach as closely in the subsurface also did not fit the data.

For the case with secondary back thrusts, I consider the scenario in which the faults intersect at depth (Figure 3.2b). Different fault geometries are modeled by varying the dip angles, dip directions and slip amount. The best fit to STRB topography is obtained with a primary fault dip angle of 15°, secondary fault dip angle of 25°, and down-dip slip of 1.5 km tapering to 0 km at the lower fault tip. The topographic profile generated from this model (Figure

3.2d) resulted in broad ridges, but the slopes are asymmetric. Some models exhibit a flat-topped morphology with a sag in the middle, and thus provide an approximate match to the observed topography. However, these models would result in large stress concentrations at the intersection of the faults, and there is no evidence to support the layer-parallel slip in mechanically weak upper horizons for the STRB as has been proposed for wrinkle ridges (Okubo and Schultz,

2004).

These results argue against simple models assuming a single-thrust fault, a pair of symmetric thrust faults, and primary and secondary back thrust fault geometries. They also suggest that STRB ridges that are bound by symmetric steep scarps are likely bound by faults on both sides. A model in which the ridges are bound by symmetric thrust faults and the vertical uplift of the fault block is accommodated at depth by viscous lower crustal deformation may be

18 able to reproduce the observed topography, but this ductile deformation cannot be represented with COULOMB.

As a simple geometric test of the possibility of symmetric thrust faults, I calculated the tectonic lithospheric thickness required for the symmetric border faults not to cross-cut each other at depth. For a ridge width of x km and dip angle of α, symmetric thrust faults facing each

19 other will cross cut each other at a depth of tan( α)( x/2) . This shows that a lithospheric thickness of <18 km would be required for a typical fault dip of 30° and ridge width of 60 km, in order to prevent the faults from crossing at depth (Figure 3.3). This constraint is relaxed to 50 km for the maximum lock-up angle for thrust faults of 60° (Scholz, 2002). Figure 3.3 shows the crossing depth for a range of ridge widths and fault dips. This limits the maximum down-dip height of the fault, which in turn will limit the maximum displacement.

The expected displacement across a fault can be calculated from the fault length using the expected displacement:length (D:L) ratio. The typical terrestrial D:L ratio is 3 ×10 -2 (Scholz,

2002). For symmetric faults at the lock-up angle 60°, the crossing depth is between 35-80 km for ridge widths of 40-90 km. This gives maximum down-dip lengths of 40-90 km. Since the constrained down-dip fault length is less than one-half of the ridge length, the D:L ratio must be applied to twice the down-dip length, yielding maximum displacements of 2.4-5.5 km.

However, for more typical thrust fault dips of 30°, the predicted displacements for faults extending to their crossing depth are only 1.4-3.1 km. This indicates that the widest ridges could have formed by uplift of a fault block between antithetic steeply dipping thrust faults. However, a number of the ridges would require either dips that are steeper than the lock-up angle or greater displacement-length ratios to form by thrust faulting. For example, the ridge in the profile in

Figure 3.1a has a height of 3 km and a typical width of 50 km, resulting in a fault crossing depth of 14-43 km for dips of 30-60°. For this ridge to form from faults extending to their crossing depth with a D:L ratio of 3 ×10 -2, a dip of 63° would be required. Alternatively, for a typical normal fault dip of 30°, a D:L ratio of 0.1 would be required.

Figure 3.3 The crossing depth for symmetric thrust faults bordering the ridges as a function of slope for different ridge widths.

In summary, many of the ridges in the Basin and Range and STRB show both symmetric and asymmetric slopes, whereas Amenthes Rupes and other lobate scarps exhibit a distinct asymmetry. The Basin and Range and the STRB scarps have similar average slope values of

~0.14. However, Basin and Range profiles have crested tops due to fluvial erosion, whereas some STRB ridges are flattened. The elevation difference across the STRB ridges also differs from the Basin and Range, but this can be understood as a result of volcanic infill from Tharsis.

Tectonic modeling results considering a variety of thrust faulting geometries have difficulty in reproducing the topography of the STRB. From a simple geometric standpoint, the symmetric slopes, high relief, and narrow widths of STRB ridges are more consistent with an origin through

Basin and Range-style extensional tectonism.

CHAPTER 4

STRAIN INDICATORS

In this section, strain calculation from the craters of the STRB and from crustal thickness models across the STRB is discussed in detail. Later, possible strain mechanisms are compared in terms of accommodation of the strain in the STRB province.

4.1 Deformed craters of the South Tharsis Ridge Belt

The STRB contains many impact craters that can help resolve the strain history of the region. The majority of undeformed craters are circular, with only a small fraction having significantly elliptical shapes (Andrews-Hanna and Zuber, 2010). Elliptical craters occur as a result of oblique impacts and ~5% of the population of small (<150 km diameter) craters are elliptical (Andrews-Hanna and Zuber, 2010; Schultz and Lutz-Garihan, 1982; Barlow, 1988;

Bottke et al., 2000). 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). Although crater analysis can reveal the strain history, it can be non-trivial in the regions where strain is highly variable both locally and regionally, as may be the case in the

STRB. This would complicate the interpretation of craters because the craters superimposed on ridges may experience little deformation even as the province as a whole extends significantly.

A number of craters within the STRB have been buried due to volcanic infilling between the ridges. The remaining ones are mostly circular to slightly elliptical in shape, with well- preserved rims (Figure 4.1). Extensional and compressional mechanisms of STRB formation may be distinguished based on the observation of craters that are elongated or shortened in the

22 direction orthogonal to the ridge axes, respectively. Aspect ratio measurements of well- preserved deformed impact craters in the region (Figure 4.1) reveal a small number of craters that are either elongated or shortened in the direction orthogonal to the ridge axes. Craters A and

B are interpreted as shortened due to the orientation of their long axes parallel to the ridge axis, with aspect ratios of 1.03 and 1.21. Crater B has a relatively higher aspect ratio, although this is complicated by the partial volcanic burial of the western rim and superposed crater on the eastern rim. Crater C is located in the southern part of 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 local trend of the STRB ridge axes suggesting lengthening. This is supported by the observation that the crater rims parallel to the long axis are discontinuous and lower in relief, consistent with extensional necking of the rims. Crater D is located on top of a ridge, and its aspect ratio is calculated as 1.09. Its long axis orientation suggests elongation. Finally, crater E clearly pre- dates 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, and its aspect ratio may be affected by the adjacent crater to the east.

This region also contains a small number of highly tectonically deformed basins of possible impact origin, such as Koval’sky basin (Figure 4.1f). If these irregular basins are of impact origin, they have undergone a significant deformation. The linear walls shown as dashed lines in Figure 4.1f suggest tectonic deformation of the basin. Due to the significant tectonism and high amount of strain, it is not possible to decipher the strain history of this basin.

Since the aspect ratio measurements do not display a clear trend of shortening or lengthening of caters orthogonal to the ridges, it is not possible to unambiguously distinguish between extensional and compressional formation mechanisms by this approach. The

23 observation of both elongation and shortening of the craters may suggest that this region has undergone several different tectonic regimes with reversing sense of strain, which has previously been suggested for Tempe Terra on Mars (Tanaka and Golombek, 1994). This would be consistent with the observation that later-formed wrinkle ridges commonly appear to nucleate from the STRB ridges. However, this analysis is also complicated by the possibility of elliptical craters forming in low-angle impacts or from near-simultaneous impacts in crater chains.

Figure 4.1 Impact craters in the STRB, in cylindrical projection. Capital letters denote the crater names. See Figure 1.2a for the locations of craters.

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. Crustal thickness models (Neumann et al., 2008) indicate that the crust comprising the STRB is thinner than that in the highlands to the west by ~2-5 km

(Figure 4.2a, b). Although the northeastern portions of the STRB overlap with the crustal thickening associated with Tharsis, the central zone of the ridge belt exhibits a thinner crust. The thinned crust is associated with the lower topography between the ridges, as would be expected from isostatically compensated crustal thinning (Figure 4.2c, d). This thinning of the crust is indicative of extensional deformation, suggesting an extensional strain of -0.03 - -0.09 for crustal thinning by ~2-5 km relative to an initial thickness of ~60 km. The maximum horizontal extension in the region is calculated as ~95 km, based on width of the ridge belt of 1050 km.

Thinning of the crust contradicts a compressional origin, and instead indicates horizontal extension of the lithosphere. Thus, both the STRB and Basin and Range display crustal thinning and a significant amount of extension. The total extension in the Basin and Range varies between

10-300% (Parsons, 1995; Hamilton and Myers, 1966; Zoback et al., 1981; Wernicke, 1992;

Sonder and Jones, 1999).This strain based on the crustal thinning will be compared with an estimate of the amount of extensional tectonic deformation based on the relief of the ridges in the next section.

4.3 Strain from horizontal tectonic deformation

If the deformation within the STRB was purely tectonic in nature, the observed relief across the ridges could be used to estimate the amount of extension. 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 this analysis, I assumed that each ridge is bordered by two symmetric faults, and then calculated the horizontal deformation generated by these faults using topographic profiles. The amount of extensional strain from this analysis will be a lower bound, since it does not include the effects of fault rotation and ductile deformation that are thought to dominate the extension in the Basin and Range province. 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) is calculated for fault dip angles d ranging from

0° to 90° as x=2 y/tan( d). The horizontal deformation from all profiles crossing each ridge was numerically integrated to calculate the change in the area. The same procedure was carried 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 STRB.

The strains calculated for normal faulting are in the range of -0.0016 to -0.0006 for fault dips ranging from 45° to 70°, respectively (Figure 4.3a). If each ridge was bordered by a single fault, the strain would be reduced by half. In order to accommodate the strain of up to -0.09 inferred from the crustal thinning, exceedingly gentle fault dips are needed (<5°). This dip is less than the observed average dips of the scarps bounding the ridges, 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 extension of the lithosphere

(between 10-300% (Parsons, 1995; Hamilton and Myers, 1966; Zoback et al., 1981; Wernicke,

1992; Sonder and Jones, 1999)) greatly exceeds the strain that would be calculated based on ridges bounded by normal faults with typical dips. I have done the same strain calculation for the Basin and Range assuming the ranges are each bordered by a single normal fault. The average dip angles are between 45° and 75° (Walsh and Watterson, 1988), resulting in tectonic strain estimates in the range of -0.014 to -0.004, respectively (Figure 4.3b), in comparison with

27 the strain estimates of 10-300% (0.1-3.0) (Parsons, 1995; Hamilton and Myers, 1966; Zoback et al., 1981; Wernicke, 1992; Sonder and Jones, 1999). This result is consistent with the interpretation that the Basin and Range extension may have been accommodated by a combination of listric normal faulting and ductile deformation (Mohapatra and Johnson, 1998).

Similarly, I suggest that a significant fraction of the extensional strain in the STRB was accommodated by ductile deformation of the crust.

If the STRB ridges are instead bounded by symmetric thrust faults, the strain is 0.003 for an optimum thrusting dip of 30°. This would result in minor crustal thickening, in conflict with the thinner crust observed in the region.

Figure 4.3 Strain versus dip angle for (a) STRB and (b) Basin and Range.

4.4 Strain from simple flexure

The amount of extensional strain in the STRB calculated from the crustal thinning of up to -0.09 is difficult to account for on a single-plate planet such as Mars. The orientations of the ridges in the STRB approximately concentric to Tharsis are not consistent with the orientation of

28 extensional membrane-flexural stresses from Tharsis loading, which should cause radial normal faulting (Banerdt et al. 1992; Banerdt and Golombek, 1989). However, the pattern of circumferential normal faulting is consistent with the expected radial extension arising from uplift of the lithosphere over large regions, for which membrane stresses dominate over flexural stresses (Banerdt et al. 1992). For Mars, membrane stresses exceed the flexural stresses for spherical harmonic degrees <8 (Turcotte et al., 1981), corresponding to half-wavelengths >1250 km. This pattern is the opposite of that expected for smaller regions of uplift, in which bending stresses dominate and radial normal faulting is expected. A best-fit circle to the STRB yields a diameter of 2730 km, confirming that it would be dominated by membrane stresses. 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 STRB relative to the size of the planet. Uplift of the lithosphere above a flattened plume head (Campbell and Griffiths, 1990), would focus the extension at the margins of the area of uplift.

However, the expected magnitude of the strain from membrane-flexural deformation of the lithosphere is low. 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 from (Turcotte and Schubert, 2002, eqn. 3.67)

y ε = , (4.1) xx 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

0.1–1.0 ×10 -5 for a lithosphere thickness of 20 km, or 0.5–5.0 ×10 -5 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 STRB (1050 km) to calculate the radius of curvature.

Uplift of 1-10 km then leads to strains ranging from 0.2-2.0x10 -5 for a lithosphere thickness of 20 km, or 0.9-9.0x10 -5 for a lithosphere thickness of 100 km. These flexural strains are compatible with the strain estimate from the ridges themselves, but both results 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.

Similar problems are encountered if the ridges are instead treated as compressional structures. The orientations of the ridges are consistent with the orientations of the younger

Hesperian wrinkle ridges in the region. However, Tharsis loading stresses alone should cause lithospheric extension in this region (Banerdt and Golombek, 1992). The formation of the wrinkle ridges requires an additional compressional stress arising from global thermal contraction and cooling (Golombek et al., 2001; Watters, 1993). 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. Global contractional strain estimates for Mars are in the range of

1.1 ×10 -2 to 2.2 ×10 -1 (Nahm and Schultz, 2011). According to thermal history models, the total accumulated contractional strain arising from the combination of global cooling and volcanic outpouring is <0.002 by the end of the Noachian, consistent with constraints from populations of

30 strike-slip faults west of Tharsis of <0.0018 (Andrews-Hanna et al., 2008), but less than that required to form the STRB by compressional tectonism. Furthermore, the progression from strike slip faulting to wrinkle ridge formation (Andrews-Hanna et al., 2008) around the

Noachian-Hesperian boundary indicates that stresses within the STRB were not yet compressional during the formation of the ridges. Thus, from the standpoint of the strain history of Mars, a compressional origin for the ridges is not favored.

4.5 Strain from craters of Western Terra Sirenum

Since Mars is a terrestrial planet without plate tectonics, the formation of the STRB by extensional tectonism must be accommodated by compression elsewhere. One candidate region to accommodate this extension is found in western Terra Sirenum. Although the craters in the

STRB did not yield a consistent sense and orientation of strain, further 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 4.4). This section quantifies the strain in western Terra Sirenum using the population of heavily deformed impact basins and craters. I later discuss the geodynamic implications of the paired extension and compression at the STRB and western Terra Sirenum.

4.5.1 Strain from Rose Diagrams

The shapes of the seven largest impact basins clearly suggest that Terra Sirenum has undergone major ductile deformation (Figure 4.4). These basins are the largest and oldest strain

31 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 understand and quantify the deformation in the Western Terra Sirenum, I measured the strain of each crater and used rose diagrams and strain ellipses to represent our results. The crater rims were traced using MOLA topography overlain with THEMIS daytime infrared data. The best-fit circles to the crater rims from a Mercator projection were found by a

Figure 4.4 Craters of Western Terra Sirenum, marked craters are the 7 largest basins with diameters >170km, in Mercator projection. Latitude – longitude range is 15 -55°S – 175-215°E

33 result in crater elongation in the direction perpendicular to the ridges, while a compressional origin should result in shortening. In order to quantify the elongation, radius of the best fitted circle ( r) is subtracted from the radius of the local crater ( R). The local elongation as a function of azimuth for each crater was normalized to the mean radius to give the relative strain as a function of azimuth, from which the rose diagrams were generated.

I analyzed 204 craters in this region with diameters greater than 10 km (Figure 4.4). In order to quantify the deformation, the maximum, minimum, and root-mean-square (RMS) strains were calculated for each crater. Positive and negative strain values represent extension and compression, respectively.

The root mean square (RMS), maximum (elongation), and minimum (shortening) strains for the seven heavily deformed basins are 0.09, 0.17, and -0.18, respectively. The RMS, maximum (elongation), and minimum (shortening) horizontal deformation for these basins are

180, 340, and -360 km, respectively. For comparison, the equivalent strain values for the craters

<170 km in diameter are 0.04, 0.091, and -0.085. Rose diagrams of the heavily deformed basins indicate that this region experienced significant deformation, dominated by NW-SE extension and/or NE-SW compression (Figure 4.6a). Individual basins display different strain directions but the average rose diagram confirms the regional trend (Figure 4.6b). The axis of maximum shortening is oriented approximately in the direction of the STRB, consistent with the possible interpretation that extension in the STRB could have been accommodated by compression in western Terra Sirenum.

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

34 shapes of the deformed basin rims, lacking linear rim segments or discrete tectonic offsets, indicate ductile rather than brittle/tectonic deformation. This differs from the morphology of the oldest impact basins in the STRB, which display tectonic deformation with linear crater rim walls (Figure 4.1f). This change in deformation style suggests a lateral transition between brittle and ductile deformation between these two regions.

Figure 4.6 Average strain Rose diagrams (a) large 7 basins (b) all craters diameters < 170 km (10 degree intervals).

4.5.2 Timing of the deformation

The timing of the deformation in the STRB is not easily constrained, since the majority of craters in the region are buried beneath the volcanic plains between the ranges and the majority of craters appear to have escaped modification due to the localized nature of the deformation.

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 using crater size- frequency distributions (Hartmann and Neukum, 2001). The age estimated for the deformed

35 basin population greater than 170 km in diameter is about ~ 4.03 Ga (Figure 4.7). I emphasize that this is a model age only, and that significant uncertainties exist in any age from crater counting, particularly those at the time of the controversial late heavy bombardment or terminal cataclysm of the inner Solar System (Gomes et al., 2005). Undeformed craters with diameters of

30-170 km yield an age of ~3.7 Ga, indicating that deformation had largely ceased by that time.

Thus, intense deformation of western Terra Sirenum began sometime during the pre-Noachian period, and continued to about the Noachian-Hesperian boundary. Possible volcanism led to the removal of a significant fraction of the small (<30 km) undeformed craters after ~3.7 Ga, as supported by a roll-over in the crater size-frequency distribution at that diameter. If the deformation in western Terra Sirenum was responsible for accommodating the strain in the SRB, it would have been synchronous with STRB development, implying a similar history for STRB tectonism.

Figure 4.7 Crater size frequency distributions of all crater population.

CHAPTER 5

THERMAL MODELING

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), a wide rift (e.g., the Basin and Range province of the western United States), or a core complex (e.g., southern Tibet). The observed mode of lithospheric extension places constraints on those parameters. I have applied a model of pure shear lithospheric extension (Buck, 1991) to the STRB in order to predict the mode of rifting for early Mars in an effort to constrain the parameters governing that deformation.

Extensional deformation of other regions within Tharsis, such as Valles Marineris, Tempe

Terrra, Alba Patera, and Coracis Fossae, were investigated previously using the same modeling techniques (Anderson and Grimm, 1998; Grott, 2005; Kiefer and Johnson 1995; Wilkins et.al.,

2002).

Lithospheric extension is controlled by the balance between the yield stress, thermal buoyancy, and compositional buoyancy. The mode of lithospheric extension is based on the strain localization or de-localization (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 results 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 a narrow rift or necking (Buck, 2007).

Necking is a self-reinforcing process since the extension becomes more rapid when the necked area is weaker. Buoyancy also affects the style of rifting. When the hot, low viscosity lower crust flows into the region of crustal stretching, the crustal thickness variations are reduced. This 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 and extension.

Wide rifting occurs when the delocalizing effects are larger than the localizing effects.

Thermal diffusion has a delocalizing effect and leads to thickening and strengthening of the brittle part of the lithosphere, which will favor wide-rifting. Extension also replaces crust with stronger mantle material, which has a delocalizing effect.

The strain rate, crustal thickness, and heat flow are the main factors influencing the style of the extension in this model. The strain rate influences the style of extensional deformation both because it affects the thermal response to rifting 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 ×10 -16 – 5×10 -15 s-1) (Sonder and Jones, 1999). In our computations, the strain rates are taken between 10 -14 and 10-17 s-1, encompassing the range of strain rates in the Basin and Range province as well as most estimates for early Mars.

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, and has been constrained to between 38 and 62 km

(Wieczorek and Zuber, 2004). The mean crustal thickness within the STRB is 55 km (Neumann et al., 2008), but this has been thinned relative to the mean value to the west of 60 km.

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 (wide and narrow rifting, and core complexes) are predicted, whereas for the wet quartz only narrow mode and core complexes are predicted (Buck, 1991). Both wet and dry diabase rheologies have been adopted for Martian crust (Grott 2005; Grott and Breuer, 2008). In our computations I have used wet diabase

(Caristan, 1982) for the crust and wet olivine (Karato et al., 1986) for the mantle.

5.1 Theory and Method

In order to investigate the thermal effects of rifting, temperature changes are modeled at the center of the extending region using the one dimensional advection-diffusion equation (Buck,

1991; Grott, 2005)

∂T ∂ 2T ∂T = κ − v , (5.1) ∂t ∂z 2 z ∂z where T is temperature, t is time, vz is the downward vertical velocity ( vz = −ε&z ;ε& being the strain rate) and κ is the thermal diffusivity. The thermal conductivity of the martian crust and mantle is thought to vary between 2.5 and 4 W m -1K-1 (Grott and Breuer, 2008). Crustal densities for southern highlands are estimated to be 2500-3000 kg m -3 (McGovern et. al., 2004), leading a thermal diffusivity of ~1 ×10 -6 m2 s-1. The parameters used in the calculations are given in Table

5.1.

In our model the lithosphere was extended to a total accumulated strain of 2.5% at strain

-17 -14 -1 rates of 10 -10 s , corresponding to rifting velocities ( vx) of 0.03 and 3 cm/yr, and time spans

(δt) of 80 and 0.8 Myr, respectively.

Table 5.1 Model Parameters of Mars.

Symbol Name Units Value

G gravity constant m/s 2 3.72

 thermal diffusivity m2/s 10 -6

3 ρc crustal density kg/m 2700

3 ρm mantle density kg/m 3500

α thermal expansion coefficient K-1 3x10 -5

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 m -1 K-1. As in previous studies, I have neglected radiogenic heat production within the crust, but this is expected to have a small effect on the results.

The advection-diffusion equation (5.1) was solved using an implicit finite difference

scheme, with a mesh size of ∆z = 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):

L dF tb = g ραT (z)zdz , (5.2) ∫0

where L is the lithospheric thickness, α the thermal expansion coefficient, ρ is the density, and

∆T denotes the temperature change at center of the rift during the extension (Buck, 1991;

Nimmo, 2004). In the computations I set ρ = ( ρc + ρm)/2. The thermal buoyancy change in Eq.

(5.2) corresponds to the change in horizontal pressure due to thermal buoyancy integrated through the vertical column, and is not strictly speaking a force. The base of the lithosphere is defined as the brittle-ductile transition (BDT) depth.

5.2 Crustal buoyancy force

The change in crustal thickness h during the extension is described by (Buck, 1991;

Grott, 2005; Nimmo 2004):

∂h ∂ 2h ∂h ∂u = κ f 2 − vx − h , (5.3) ∂t ∂x ∂x ∂x

where κ f is the effective flow diffusivity and vx is the applied horizontal rifting velocity. The

rifting velocity vx varies linearly between zero at the rift center and ε&X e at the end of extension zone x = X e, where the zone of extension is between -Xe and Xe. Eq. (4) is discretized and solved

using an implicit finite difference scheme between x=0 and XL, where XL is the half-width of a

∂h ∂h region of initially undeformed lithosphere. The boundary conditions are | = 0, | = 0, ∂x x=0 ∂x X L

with Xe = 40 km and XL = 250 km. The effective flow diffusivity κ f depends on time and it is given by (Buck, 1991):

3 gρ *  hRT 2   − E  M   κ f =   exp   (5.4) C  E(TM −TS )  RT M 

where C = A-110 6 Pa 1-n is the pre-exponential term in the effective viscosity relation (Buck,

* 1991), ∆ρ = ( ρm -ρc ) ρc /ρm, where ρc and ρm are the crustal and mantle densities, respectively, and TM is the Moho temperature.

The change in crustal buoyancy force is computed by

h ρc dF ch = pdz ≈ gh h(ρm − ρc ) , (5.5) ∫0 ρm

where h denotes the crustal thickness change at the middle of the rift during the extension.

5.3 Yield strength

The yield strength of the crust is given by

h F = σ (T, z,ε )dz (5.6) ys ∫0 &

where σσσ is the lesser of the frictional strength or the ductile strength. The frictional stress σ b for extension according to Byerlee’s law as given in (Grott and Breuer, 2008) by neglecting the yield stress, and only considering the frictional strength on already-failed planes:

0.786 σ v , σ v ≤ 529.9 MPa σ b =  (5.7) 56.7 MPa + 0.679 σ v , σ v > 529.9 MPa

where σ v = βρ r gz . is the lithostatic pressure (Kohlsteadt et. al., 1995). I have scaled the lithostatic pressure with a factor of β=0.6 to account for the reduction in stress due to both horizontal tension and hydrostatic pore pressure.

The ductile flow stress (the stress to cause ductile deformation at a specific strain rate and temperature) is:

1/ n  ε&   E  σ d =   exp  , (5.8)  A   nRT 

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 5.2.

Material n A E C = A-1(1.0x10 6 Pa) 1-n

(Pa -ns-1) (KJ mol -1) (Pa-s)

Pyroxene (1) 5.3 2.5 ×10 -37 380 7.3x10 10

Wet diabase (2) 3.05 3.1x10 -20 276 1.6 ×10 7

Dry diabase (2) 4.7 1.1x10 -26 488 5.7 ×10 3

Dry olivine (2) 3.5 2.4 x10 -16 540

Wet olivine (2) 3.0 1.9x10 -15 420

5.4 Results

I computed the change in the total force balance from equations 5.2, 5.5, and 5.6 for a

range of thermal gradients, crustal thicknesses, and strain rates (Figure 5.1). I first examine the

predicted mode of extension for two representative strain rates (Figure 5.1a, b). I 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 slower strain rates and corresponding smaller rifting

velocities, wide rifts occur for a larger range of thermal gradients and crustal thickness (Figure

5.1a, b).

I 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). For a typical Basin and Range

strain rate of a 5 ×10 -15 s-1, total force is positive for the expected Martian thermal gradients

(Figure 5.1c). Positive total force indicates that de-localization effects are greater than the localization effects, and thus predicts wide rifts. Strain rates of 10 -17 -10 -14 s-1 are 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

>20 K/km (Schultz and Watters, 2001; McGovern et al., 2004; Grott, 2005), wide rifting can occur for all strain rates >10 -17 s-1, whereas narrow rifting can only occur for strain rates >10 -15 s-

1. Increasing the strain rate shifts the wide to narrow rift transition to higher thermal gradients for a given crustal thickness (Figure 5.1d). For the expected thermal gradients (>15 K/km) and typical Basin and Range strain rates (2 ×10 -16 – 5×10 -15 s-1; Sonder and Jones, 1999), wide rifting

or core complex formation is preferred for all but the highest strain rate and lowest thermal

gradient (Figure 5.1d).

An approximate estimate of the strain rate in the STRB can be made based on the

constraints on the timing and total strain. Since the formation of the STRB is constrained to be

before the Noachian-Hesperian transition at ~3.7 Ga, based on the timing of deformation in

western Terra Sirenum and the presence of Hesperian plains between the STRB ridges, 800 Myr

is an upper bound on the duration of extension. Extension likely occurred after the late heavy

bombardment, suggesting an upper bound of 200 Myr. If I take the total strain as 0.09 (from the

crustal thinning) or 0.18 (from western Terra Sirenum craters), the lower bound on the strain rate

ranges between 5.7x10 -18 to 1.4x10 -17 for a time period of 500 Myr. An upper bound for the

strain rates can be obtained by taking the typical tectonic velocities on Earth (~1 cm/yr) and

44 dividing by the total width of the STRB (1050 km), yielding a strain rate of 3.0 ×10 -16 s-1. For the estimated thermal gradients of Noachian Mars, these strain rates predict either wide rift or core complex mode extension.

Figure 5.1 Rifting mode diagrams for strain rate 8x10 -15 s-1 (a) 1x10 -16 s-1 (b). Total forces for a 60 km thick crust and the strain rate 5 ×10 -15 s-1s-1 (c) and semi-log rifting mode diagram (d) for a 60 km thick crust. The range of thermal gradients estimated for the Noachian period is indicated by the shading in (c) and the range of strain rates estimated for the Basin and Range (2 ×10 -16 – 5×10 -15 s-1) is indicated by the shading in (d).

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. The range of the thermal gradients and crustal thicknesses 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 plateau. Alternatively, narrow rifting at Coracis Fossae has been explained by invoking a thinner crust at the time of extension

(Grott, 2005), requiring a later stage of crustal thickening after the formation of rifts. 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,

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.

CHAPTER 6

DISCUSSION AND CONCLUSIONS

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. Our results favor an extensional origin, but some of the observations are ambiguous. In this section, I discuss the significance of both compressional and extensional interpretations.

6.1 Compressional origin

A compressional origin for the STRB may be supported by the morphology of the ridges

(Schultz and Tanaka, 1994) and the observation of shortening of some craters in the region.

However, this would require a significant source of compressional stress in the direction radial to

Tharsis as discussed in Section 4.4. The observed thinner crust within the STRB is the main argument against the compressional model. To explain this observation in a compressional scenario, it would be necessary to invoke a thinner crust prior to the formation of the ridge belt, generated by an independent and unknown mechanism. This pre-existing thinner crust would result in a stronger lithosphere, due to the coupling of the mantle and crustal lithosphere (Grott and Breuer, 2008), which could impede the formation of the ridge belt by compression.

However, as discussed earlier, for the case where the STRB is bounded by two thrust faults dipping towards each other, the thickness of the lithosphere should be less than 20 km for an average ridge width of 60 km. Thus, a compressional origin for the STRB is incompatible with the crustal thinning in the region, and would also require an unknown source of large

47 compressional stress to produce the radial compression in STRB. Therefore, a compressional mechanism is not favored, though it cannot be ruled out.

6.2 Extensional origin

An extensional mechanism for the STRB is favored based on the morphologic and topographic similarity of the ridges to the Basin and Range province. This mechanism also explains the thinner crust observed in STRB. Furthermore, wide rift extension as in the Basin and Range should be the favored mode of extension in the southern highlands of Mars. An extensional origin leads to different implications for early Martian geodynamics.

Previous studies have suggested that Tharsis may have formed in response to a giant spherical harmonic degree-one mantle plume (Harder and Christiansen, 1996). The early stages of Tharsis formation may thus have involved plume-induced uplift of the lithosphere, before the rise transitioned into a flexurally supported load, leading to an evolving stress regime in the area.

In the case of a plume-uplift for Tharsis, membrane stresses will dominate causing radial extension within the area of uplift (Banerdt et. al., 1992). 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. 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 grabens observed in the area. Then,

Hesperian-aged circumferential wrinkle ridges would develop as a result of significant global

cooling and contraction (Golombek et al., 2001; Watters, 1993). The transition to compression

could also allow reactivation of normal faults in a thrusting sense, making it difficult to interpret

the tectonic style of the ridge belt. This evolution may help to explain the presence of both

elongated and shortened craters in the direction orthogonal to the ridge axes.

However, the biggest challenge to an extensional origin is to generate the inferred strain.

The strain of ~0.09 inferred from the crustal thinning is incompatible with the expected

magnitude of strain from plume-induced uplift. Strains from membrane and flexural

deformation are generally low (of order 1%), and cannot account for significant net extension of

the lithosphere as discussion in Section 4.4. Crustal thinning by simple tectonic extension would

require exceedingly low fault dip-angles (<5°), in conflict with the dip angles suggested for the

Basin and Range of 45°-60° (Hamilton, 2005). However, wide rifting prefers a ductile, hotter

crust (Buck, 1991), which suggests the importance of ductile deformation in Basin and Range.

Thus, crustal thinning may have occurred by a combination of tectonic extension of the upper

crust and ductile necking of the lower crust.

It is also necessary to accommodate the extensional strain in the STRB with compression elsewhere, since Mars doesn’t have plate tectonics. The craters of western Terra Sirenum display significant ductile deformation, comparable in magnitude to the extension in the STRB.

However, the crust of Terra Sirenum is not thickened as would be expected from compression.

In the next section, we propose a scenario linking the tectonic evolution of Tharsis, the STRB, and western Terra Sirenum.

6.3 Proposed tectonic and geodynamic evolution

Wide-rift style tectonic extension of the STRB may have been triggered by plume- induced uplift of Tharsis, which would create 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.

Recently, parameterized convection models (Morschauser, et al., 2011) have shown that the 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. The ductile deformation observed in these basins may have been encouraged by the thicker crust and higher heat flow prior to lower crustal recycling. This crustal shortening and recycling in Western Terra Sirenum would have enabled wide rift style extension to continue in the STRB, accumulating strains of -0.1 and +0.1 in these regions, respectively. Later, the radial extension within STRB would be replaced by circumferential extension due to the transition to membrane-flexural support of Tharsis, forming dike-induced radial graben. Gradual cooling and contraction of Mars during Hesperian would have increased the horizontal compressional stresses leading to the formation of circumferential thrust faults due to radial shortening within the STRB.

I emphasize that the above scenario is speculative. An extensional origin for the STRB cannot be confirmed definitely by our geological and geophysical analyses. The nature and cause of the significant ductile deformation in the Western Sirenum is unclear. Furthermore, the

50 geodynamic mechanism responsible for Tharsis formation is uncertain. Nevertheless, many aspects of the tectonic and geodynamic evolution of 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 there is no evidence for plate tectonics having ever operated on early

Mars, the possibility of significant horizontal strains, ductile deformation, and even lower crustal recycling cannot be ruled out. 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 semi-mobile lithosphere.

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APPENDIX A

SLOPE CALCULATION

Slopes were measured from the steeply sloping sides of the ridges, avoiding locally over- steepened areas to obtain a more representative value in each side (western and eastern) of the profile. Later, average slope for each ridge is calculated by averaging the slopes from related profiles. Final slope values for each region are obtained from averaging the average slopes of the ridges within the region. Table 3.1 shows this average value, and profiles used in this calculation is given in the rest of the Appendix A. Representative profiles (Figure 3.1, and red profiles in Figures A-1, A-3 and A-5) are selected from less cratered, deformed and eroded parts of the ridges.

South Tharsis Ridge Belt

Locations of the profiles:

Ridge A:

Ridge B:

Ridge C:

Figure A-2. Individual profiles of the South Tharsis Ridge Belt plotted all together (left) and separately (right). The vertical exaggeration is 20× in all panels. See figure A-1 for the locations of the profiles. Please note the variable scales within the profiles.

Ridge D:

Ridge E:

Ridge F:

Figure A-2. (continued) Individual profiles of the South Tharsis Ridge Belt plotted all together (left) and separately (right). The vertical exaggeration is 20x in all panels. See figure A-1 for the locations of the profiles. Please note the variable scales within the profiles.

Basin and Range:

Locations of individual profiles:

Ridge A:

Ridge B:

Ridge C:

Figure A-4. Individual profiles of the Basin and Range plotted separately (left) and all together (right). The vertical exaggeration is 20 × in all panels. See figure A-3 for the locations of the profiles. Please note the variable scales within the profiles.

Ridge D:

Ridge E:

Ridge F:

Figure A-4. (continued) Individual profiles of the Basin and Range plotted all together (left) and separately (right). The vertical exaggeration is 20 × in all panels. See figure A-3 for the locations of the profiles. Please note the variable scales within the profiles.

Ridge G:

Ridge H:

Ridge I:

Figure A-4. (continued) Individual profiles of the Basin and Range plotted all together (left) and separately (right). The vertical exaggeration is 20x in all panels. See figure A-3 for the locations of the profiles. Please note the variable scales within the profiles.

Ridge J:

Ridge K:

Ridge L:

Figure A-4 (continued) Individual profiles of the Basin and Range plotted all together (left) and separately (right). The vertical exaggeration is 20x in all panels. See figure A-3 for the locations of the profiles. Please note the variable scales within the profiles.

Ridge M:

Ridge N:

Ridge O:

Amenthes Rupes

Locations of individual profiles:

Figure A-6. Individual profiles of the Amenthes Rupes plotted all together (left) and separately (right). The vertical exaggeration is 20 × in all panels. See figure A-5 for the locations of the profiles. Please note the variable scales within the profiles.

APPENDIX B

THERMAL MODELING

Thermal modeling assumptions:

For the tectonic extension, the model of pure shear extension of McKenzie (1978) and

England (1983) is used as in Buck (1991). Initial lateral crustal thickness variations are not taken into account and heat transport in the lateral direction is neglected. Buck’s (1991) model of lithospheric extension is based on one-dimensional heat transport in the vertical direction. Also an environment with a constant thermal gradient at the base of the lithosphere as a function of space is considered, following Buck (1991) and Grott (2005). The surface heat fluxes for diverse regions on Mars have been estimated from the effective elastic lithosphere thickness (Grott et al.,

2005, Ruiz et al., 2008) and from the brittle-ductile transition (BDT) depths beneath large thrust faults (Schultz and Watters, 2001; Grott et al., 2007; Ruiz et al., 2008). The thermal gradients are computed using the Fourier’s law q=−k dT/dz, where the thermal conductivity of the crust is taken to be k=2.5Wm −1K−1. The crustal radiogenic heat production appearing on the right hand side of equation (3) (Buck, 1991) is neglected in the thermal modeling as in Grott (2005) because it is very small. A single layer brittle–ductile crust overlying a single layer brittle-ductile mantle is used and the densities of the crust and mantle are assumed to be constant, as in Buck’s methodology (1991).

Figure B-1 shows the effect of different mantle rheologies on the depth to the brittle- ductile transition (BDT) as a function of the thermal gradients. For the crust, wet diabase is used for the final models. Wet and dry olivine and pyroxene show a significant difference only for low thermal gradients. For wet diabase, the BDT depth becomes smaller with increasing thermal gradients. This implies that for thinned crust the ductile behavior will be a dominating factor.

The BDT points are lower for materials like pyroxene and clinopyroxene with stronger rheologies. There is a little difference between the BDT points for wet and dry olivine used for the mantle. For Mars, mostly wet and dry/wet olivine rheologies are used, while clinopyroxene and dry olivine are mostly used for Venus.

Figure B-1. Brittle-ductile depth change for different rheologies.