Select problems in planetary structural Geology: Global-scale tectonics on Io, regional-scale kinematics on Venus, and local-scale field analyses on Earth with application to Mars

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Authors Jaeger, Windy Lee

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Link to Item http://hdl.handle.net/10150/280768 SELECT PROBLEMS IN PLANETARY STRUCTURAL GEOLOGY: GLOBAL-

SCALE TECTONICS ON 10, REGIONAL-SCALE KINEMATICS ON VENUS, AND

LOCAL-SCALE FIELD ANALYSES ON EARTH WITH APPLICATION TO MARS

by

Windy L. Jaeger

Copyright © Windy L. Jaeger 2005

A Dissertation Submitted to the Facuhy of the

DEPARTMENT OF PLANETARY SCIENCES

In Partial Fulfillment of the Requirements For the Degree of

DOCTOR OF PHILOSOPHY

In the Graduate College

THE UNIVERSITY OF ARIZONA

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Copyright 2005 by Jaeger, Windy L.

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As members of the Final Examination Committee, we certify that we have read the dissertation prepared by Windy L. Jaeger entitled Select Problems in Planetary Structural Geology: Global-Scale Tectonics on

lo. Regional-Scale Kinematics on Venus, and Local-Scale Field Analyses on

Earth with Application to Mars

and recommend that it be accepted as fulfilling the dissertation requirement for the

Degree of Doctor of Philosophy

ixhloif. Victor R. Baker date / > / y date

Alfred S. McEwen date

Robert G. Strom date

Elizabeth PT Turtle date

Final approval and acceptance of this dissertation is contingent upon the candidate's submission of the final copies of the dissertation to the Graduate College. I hereby certify that I have read this dissertation prepared vmder my direction and recominend that it bs.accepted as fulfilling the dissertation requirement.

Dissertation Ehrector: Victor R. Baker date 3

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SIGNED: 4 ACKNOWLEDGEMENTS

This dissertation is not only a compilation of my research, but a reflection of the invaluable contributions and support I received from a great number of people and institutions. The financial support that made this work possible was provided by NASA's Galileo Project, NASA Planetary Geology and Geophysics Program grants NAG5-10166 and NAGW5-3632, and Interagency Agreement W-10,265. I would first like to thank my advisor, Vic Baker. It is a privilege to have had such a fine mentor. I am also gratefiil to the former teachers and supervisors who shaped my interests in geology and planetary science, namely Karl Karlstrom, Horton Newsom, Mike Drake, Ricardo Schwarz and Mr. Max. Much of the research described in these pages was collaborative work intended for journal publication, and I am truly indebted to the following people for their roles as coauthors on those manuscripts: Zibi Turtle, L^zlo Keszthelyi, Devon Burr, Josh Emery, Alfred McEwen, Vic Baker, Jani Radebaugh, Bob Pappalardo and Hideaki Miyamoto. The content of this dissertation was also greatly improved by suggestions from my committee members: Vic Baker, Paul Kapp, Alfred McEwen, Bob Sfrom and Zibi Turtle. I thank Bob Strom and Ken Tanaka for help in working with Magellan data, Mary Chapman and Magnus Gu6mundsson for participating in a day of perilous fieldwork, Mark Orsen for kindly supplying me with his personal file on the Odessa Craters, and my fellow lo researchers (Alfred McEwen, Laszlo Keszthelyi, Zibi Turtle, Moses Milazzo, Jani Radebaugh, Ashley Davies, Julie Rathbun, Lionel Wilson, John Spencer, Paul Geissler, Dave Williams, Bob Howell, Paul Schenk, Jason Perry, Rosaly Lopes...) for being the most congenial working group in planetary science. I'd also like to extend my appreciation to the following fiiends who made even the hottest Tucson days pleasant: Gwen and Jason Barnes, Zibi Turtle, Ralph Lorenz, Joe Spitale, Eric Jensen, Ingrid Daubar, Rachel Masfrapa, Jen Grier, Andy Rivkin, Nancy Chabot, Jani Radebaugh, Barb Cohen, Moses Milazzo, Kerry Granger, Paul Withers, Fred Ciesla, Chris Schaller, Pete Lanagan, Dave O'Brien, Ross Beyer, Terry Hurford, Jonathan Fortney, and Betty Pierazzo. I am especially grateful to Tina Dajic and Tamara Legendre for their long-lived friendships. Tina's enthusiasm for life is always refreshing, and her compassion for others seems endless. Tamara has been my partner in crime since age 5, and what more can I say? She was my childhood fiiend, my gymnastics teammate, my undergraduate roommate, my road trip/camping buddy... Finally, I would like to thank my family. Words cannot adequately express my gratitude for the love and support they've given me through the years. The combination of my father's wisdom and my mother's creativity fostered in me the curiosity and ingenuity of a scientist. My parents also imparted to me their love for the outdoors and nurtured my somewhat peculiar fascination with rocks. I am a geologist because of them. I sincerely thank my sister, Monica, who has been and will always be my best fiiend, and her children, Garrett and Raquel, who supply much-needed comic relief and whom, despite Copernicus' claim, my world revolves around. Above all, I extend my deepest appreciation to my husband, Laszlo, for his unwavering patience, kindness and love. He is my colleague, my friend, my proofreader, my companion, my field partner, my love. DEDICATION

To my father, for instilling in me a love of science and nature. 6

TABLE OF CONTENTS

LIST OF FIGURES 9

LIST OF TABLES 11

ABSTRACT 12

CHAPTER 1: INTRODUCTION 14

1.1: Overview of lo 14

1.2: Overview of Venus 18

1.3: Overview of the Channeled Scabland in Washington State 21

1.4: Overview of Athabasca Valles, Mars 25

CHAPTER 2: OROGENIC TECTONISM ON lO 31

2.1: Introduction 32

2.2: A Link Between Tectonism and Volcanism on lo 34

2.3: Orogenic Tectonism on lo 51

2.3.1: Analysis of Previous Models 51

2.3.2: Estimating the Thickness oflo's Lithosphere 57

2.3.3: Can Hotspot Swells Influence Ionian Mountain Formation? 65

2.4: Observed Structures 69

2.4.1: Mountain Morphology 69

2.4.2: Kinematic Indicators 72

2.4.3: Crosscutting Relationships 75

2.4.4: Summary of Observations 76 7

TABLE OF CONTENTS - continued

2.5: Conclusions 78

CHAPTER 3: A KINEMATIC ANALYSIS OF NEFERTITI CORONA, VENUS..80

3.1: Kinematic Indicators at Nefertiti Corona 81

3.1.1: Location A 81

3.1.2: Location B 85

3.1.3: Location C. 87

3.1.4: Location D 89

3.1.5: Location E 92

3.1.6: Location F. 92

3.2: Kinematic History of Nefertiti Corona 95

3.3: Implications 96

3.4 Conclusions 100

CHAPTER 4: BASALTIC RING STRUCTURES IN THE CHANNELED

SCABLAND: IMPLICATIONS FOR MARS 101

4.1: Introduction 101

4.2: Field Observations of the Odessa Craters 102

4.3: Banks Lake Cross Section 124

4.4: The Origin of the Odessa Craters 127

4.5: Basaltic Ring Structures Near Tokio Station, WA 130

4.6: Comparison to Features in Athabasca Valles, Mars 132 8

TABLE OF CONTENTS - continued

4.7: Conclusions 136

CHAPTERS: SUMMARY 138

REFERENCES 140 9 LIST OF FIGURES

FIGURE 1.1, False-color hemispheric view of lo 15

FIGURE 1.2, Hemispheric view of the surface of Venus 20

FIGURE 1.3, Maps of the CRBG and the Channeled Scabland 22

FIGURE 1.4, MOLA topography over Athabasca Valles 26

FIGURE 1.5, Geomorphic/thermophysical map of Athabasca Valles 28

FIGURE 2.1, Ionian paterae 36

FIGURE 2.2, Locations of ionian mountains 43

FIGURE 2.3, Shamshu Patera 45

FIGURE 2.4, Ionian volcanoes 46

FIGURE 2.5, Tectonic mountains on lo 47

FIGURE 2.6, Mountain-patera association 49

FIGURE 2.7, Predicted stress state in lo's lithosphere 53

FIGURE 2.8, Volumetric strain due to subsidence stress and thermal stress 60

FIGURE 2.9, Net effect of subsidence and thermal stresses in the ionian lithosphere 63

FIGURE 2.10, Predicted surface deformation from an ionian hotspot swell 67

FIGURE 2.11, Monan Mons 73

FIGURE 2.12, Tohil Mons 77

FIGURE 3.1, Magellan radar and altimetry data over Nefertiti Corona 82

FIGURE 3.2, Locations of Figures 3.3-3.8 83

FIGURE 3.3, Location A 84 10 LIST OF FIGURES - continued

FIGURE 3.4, Location B 86

FIGURE 3.5, Location C 88

FIGURE 3.6, Location D 90

FIGURE 3.7, Location E 93

FIGURE 3.8, Location F 94

FIGURE 3.9, Kinematic history of Nefertiti Corona 97

FIGURE 4.1, Map of field sites in eastern Wahsington 103

FIGURE 4.2, Airphotos of Odessa Craters 105

FIGURE 4.3, Field map of lithofacies at Tule Crater 107

FIGURE 4.4, Field sites at Amphitheater Crater 109

FIGURE 4.5, Field sites at Cinnamon Roll Crater 111

FIGURE 4.6, Field sites at Colosseum Crater 113

FIGURE 4.7, Field sites at Rock Rose Crater 116

FIGURE 4.8, Field sites at Tule Crater 118

FIGURE 4.9, Structural data at Colosseum Crater 122

FIGURE 4.10, Stereonet plots of structural data at Cinnamon Roll Crater 123

FIGURE 4.11, The Banks Lake structure 125

FIGURE 4.12, Models for the formation of the Odessa Craters 128

FIGURE 4.13, Field map of lithofacies present in the Tokio BRSs 131

FIGURE 4.14, Comparison of Tokio BRSs and martian ring structures 133 11 LIST OF TABLES

TABLE 2.1, Catalog of ionian mountains 38

TABLE 2.2, Statistics on ionian mountains 48

TABLE 4.1, Locations and sizes of the five Odessa Craters examined in this study 106

TABLE 4.2, Trends and plunges of basalt columns at Amphitheater Crater 110

TABLE 4.3, Trends and plunges of basalt columns at Ciimamon Roll Crater 112

TABLE 4.4, Trends and plunges of basalt columns at Colosseum Crater 114

TABLE 4.5, Trends and plunges of basalt columns at Rock Rose Crater 117

TABLE 4.6, Trends and plunges of basalt columns at Tule Crater 119

TABLE 4.7, Comparison of BRSs in the Channeled Scabland and MRSs 135 12

ABSTRACT

lo's mountains are cataloged in order to investigate their formation. Of the 101 mountains imaged with sufficient coverage and resolution for further analysis, 4 are volcanoes, and 97 are tectonic massifs. Of the 97 tectonic mountains, >40 abut paterae

(volcanic or volcano-tectonic depressions). This juxtaposition is unlikely to be coincidental as the probability of it occurring by chance is -1.08%. The observed mountain-patera association may be due to orogenic faults acting as conduits for magma ascent, thus fueling patera formation near mountains.

As resurfacing buries a shell of material fi-om lo's surface to the base of the lithosphere, its effective radius is reduced and it heats up. The volume change due to subsidence and thermal expansion is calculated as a function of lithospheric thickness.

Conservation of volume dictates that this material is uplifted at lo's surface. By estimating the total volume of mountains, lo's average lithospheric thickness is constrained to >12 km.

A kinematic analysis of Nefertiti Corona, Venus, reveals that the corona's interior moved east as a relatively coherent thrust sheet with most deformation occurring on the distal margin. Additionally, an en-echelon array indicates a history of semi-brittle deformation on the northern side of Nefertiti's tectonic annulus. Regional heating from the thermal diapir that formed Nefertiti probably reduced the crustal viscosity and enabled the semi-brittle deformation. 13

The "Odessa Craters" in the Channeled Scabland of eastern Washington State are basaltic ring structures (BRSs) 50-500 m in diameter that are comprised of discontinuous, concentric outcrops of sub vertically-jointed basalt and autointrusive basaltic dikes. It is postulated that they formed as follows: phreatovolcanic activity disrupted a relatively thin, active lava flow forming rootless cones; the lava flow inflated around the cones; tensile stresses caused concentric fracturing; dikes exploited the fi-actures and fed lava to the surface; and subsequent erosive floods excavated the structures. A second population of BRSs near Tokio Station, WA, are morphologically analogous to quasi-circular structures in Athabasca Valles, Mars (a region that is geologically similar to the

Channeled Scabland). If the martian features formed as BRSs, then they indicate local water-lava interactions and at least two floods through Athabasca Valles. 14

CHAPTER 1: INTRODUCTION

Structural geology is "the study of the architecture of the Earth's crust, insofar as it has resulted from deformation" [Davis, 1984; after Billings, 1972]. A fundamental approach of structural geology is to examine small-scale features within a system to infer the large-scale kinematic history of that system. Until recently the extent to which this technique could be applied to problems in planetary science was quite limited. However, advances in spacecraft instrumentation (e.g., higher-resolution imaging) have made comprehensive studies in planetary structural geology possible. This dissertation includes three such investigations into (1) orogenic tectonism on Jupiter's moon lo; (2) the kinematic history of Nefertiti Corona, Venus; and (3) basaltic ring structures in the

Channeled Scabland with application to Mars. Collectively, these three studies broaden our understanding of the diverse nature of crustal deformation processes on planetary bodies.

1.1: Overview of lo

In 1610 Galileo Galilei peered through a primitive Earth-based telescope and discovered four moons orbiting Jupiter [Galilei, 1610]. lo is the irmermost of these

Galilean satellites (Fig. 1.1). With a mean radius of 1821 km, it is roughly the size of

Earth's Moon, and, like Earth's Moon, lo is tidally locked such that its same face points toward Jupiter at all times. Jo's density and moment of inertia indicate that it is a differentiated body with a large (-50% of the radius of lo) metallic core [e.g., Anderson 15

m-

•#

Figure 1.1. False color, hemispheric view of lo with north to the top. This image merges Galileo color data with higher resolution (up to 1.3 km/pix) SSI imaging. 16 et al., 2001]. Spectral data show that the colorful nature of lo's surface is due to a combination of sulfurous volatiles (primarily SO2 and S) and silicates [e.g., Nash et al.,

1986; Spencer and Schneider, 1996; McEwen et al., 1998; Geissler et al, 1999, 2004].

Morphologically, the surface of lo consists of vast plains (with occasional km-high scarps) disrupted by isolated, rugged mountains, paterae (i.e., volcanic ± tectonic depressions), low-relief volcanic centers, and lava flows [e.g., Masursky et al., 1979;

Schaber, 1982; Carr et al., 1998].

lo is tidally heated because its Laplace resonance with Europa and Ganymede forces it into an elliptical orbit [Peale et al., 1979; Greenberg, 1982]. As a result, it is the most volcanically active body in the solar system [e.g., McEwen et al., 1998]. However, prominent volcanic edifices are relatively uncommon on lo. Instead, lavas typically erupt from paterae and topographically indistinct volcanic centers [e.g., Schaber, 1982;

McEwen et al., 1998]. While the average surface temperature of lo is only -100 K

[Spencer and Schneider, 1996], eruptive temperatures derived from the Galileo Near

Infrared Mapping Spectrometer (NIMS) and SSI data are quite hot (with reported temperatures as high as 1870 K) and suggest mafic to ultramafic volcanism on lo [e.g.,

McEwen et al., 2000; Davies et al., 2001; Milazzo et al., in review]. Despite lo's copious volcanism and its apparent lack of plate tectonics, the vast majority of ionian mountains appear tectonic, rather than volcanic, in origin [e.g., Carr et al., 1998; Schenk et al.,

2001], and some have morphologies consistent with uplift by thrust faulting [Schenk and

Bulmer, 1998]. These mountains can reach soaring heights of up to -18 km [Schenk et al., 2001]. It is thought that the transfer of material from lo's interior to its exterior via 17 igneous processes drives vertical subsidence of the lithosphere and, in doing so, generates a large horizontal compressive stress. This subsidence stress may drive orogenic tectonism on lo [Schenk and Buhner, 1998].

A variety of data sets have been used to study lo, including ground-based telescopes and the Hubble Space Telescope, in addition to the Voyager 1, Voyager 2, Galileo and

Cassini spacecraft [e.g.. Spencer and Schneider, 1996; Turtle et al., 2004], The Voyager and Galileo images are of particular importance for the work reported here. The Voyager spacecraft were both launched in 1977 and arrived at Jupiter four months apart in 1979.

The Voyager Imaging Science Subsystem (ISS) consisted of two 800 x 800 pixel vidicon cameras, a wide-angle camera and a narrow-angle camera, each equipped with 8 filters sensitive to wavelengths fi'om ultraviolet to orange. The highest resolution images of lo

(500 m/pixel) were collected by Voyager 1. The Galileo spacecraft was launched in

October, 1989, and reached the jovian system in December of 1995. The Galileo Solid

State Imager (SSI) was an 800 x 800 pixel framing charge-coupled device (CCD) camera.

It had an 8-position filter wheel and was sensitive to violet to short infrared wavelengths.

SSI images provide hemispheric coverage of the antijovian face of lo at better than 2 km/pixel, and the highest resolution image obtained by SSI is 9 m/pixel. While the more recent Galileo SSI images are generally superior to the Voyager images, the subjovian face of lo remains best imaged by Voyager. 18

1.2: Overview of Venus

Venus is Earth's sister planet in terms of both size (Rvenus = 6052 km) and proximity

(0.28 AU at closest approach). Earth and Venus are also similar in bulk density (pvenus =

5250 kg/m^) and, because they formed in the same region of the solar nebula, they presumably accreted from similar materials. However, striking differences demarcate the two planets. Venus spins on its axis in a slow, retrograde manner such that its day (~243

Earth days) is longer than its year (-225 Earth days), and, unlike the Earth, Venus does not have a measurable magnetic field. Additionally, Venus has a thick C02-dominated atmosphere that is effectively devoid of water vapor. Because its atmosphere is so thick, the atmospheric pressure at the surface of Venus is about 92 bars, which is equivalent to the pressure at a depth of nearly 1 km in Earth's oceans. Venus has also experienced a runaway greenhouse effect, and, as a result, its surface temperature is quite high (-750

K). Beneath the thick, hot atmosphere, the surface of Venus appears geologically complex in that it has been shaped primarily by volcanic and tectonic processes. The average age of the surface, as derived from impact crater size distributions, is 190-750

Ma [Schaber et al., 1992; Strom et al., 1994; McKirmon et al., 1997], which is, to first order, similar to Earth's average surface age of -200 Ma. The spatially and hypsometrically random crater distribution on Venus as well as the paucity of lava- embayed craters and the nonrandom distribution of volcanoes suggest that the planet experienced a global resurfacing event a few hundred million years ago [Strom et al.,

1994]. However, other researchers favor a uniformitarian interpretation of the data and suggest that the resurfacing of Venus is an equilibrium process [e.g., Phillips et al., 1992]. 19

While Venus was the target of the first interplanetary probe, the Soviet spacecraft

Venera-1 in 1961, and has been revisited by several subsequent missions, the work here reUes almost exclusively on data returned by the Magellan spacecraft. Magellan was launched in 1989 and inserted into a near-polar orbit around Venus in 1990. The

Magellan synthetic aperture radar (SAR) instrument provided 100 m/pixel images of

98% of the surface of Venus. Special care must be used when making geologic interpretations based on these radar images [Tanaka, 1994], In general, the brightness of a pixel is associated with the average cm-scale roughness within the pixel. However, the orientation of slopes also has a strong affect on radar brightness. Because of the side- looking nature of the SAR, slopes modulate the radar echo strength such that those facing the antenna appear brighter and those facing away return a weaker signal and appear darker. For the same reason topography appears distorted in SAR data. Slopes that face the radar anterma are foreshortened. This has the effect of displacing peaks towards the radar look-direction. In cases where a slope facing the radar antenna exceeds the incidence angle (17° (polar)-45.7° (equatorial)), its top and bottom appear transposed in the SAR data.

The Magellan spacecraft also obtained altimetry data with a vertical resolution of 80 m and a footprint ranging from 8x11 km (equatorial) to 29 x 29 km (polar) [Tanaka,

1994]. From this dataset, a global topographic map with a resolution of ~5 km/pixel was produced [Ford and Pettengill, 1992]. The most profound result from the altimetry data is that, imlike the Earth, Venus has a unimodal hypsometry with roughly 80% of the surface having an elevation within one kilometer of the mean planetary radius (Fig. 1.2). 20

/

•IJ

1 'IniK l.il} k.ulniN I kill ) (,n4s (iii'Nf) <<>*'"^4 <>nS{> '>(!<•)<• W"

Figure 1.2. A hemispheric view of the surface of Venus centered at 0° east longitude with north to the top. This image depicts Magellan radar data and is color-coded to represent elevation. In addition to collecting radar images and altimetry data, Magellan also obtained near- global gravity field data. From the gravity data, the thickness of the venusian elastic lithosphere was estimated to be about 15-45 km over a wide range of geographic locations [Sandwell and Schubert, 1992; Smrekar, 1994; Phillips, 1994]. These values are comparable to the elastic thickness of Earth's oceanic lithosphere, which can be as large as 50 km [Caldwell and Turcotte, 1979].

1.3: Overview of the Channeled Scabland in Washington State

The Channeled Scabland in east-central Washington State is a region that has been shaped by both flood basalt volcanism and catastrophic flooding (Fig. 1.3). The primary bedrock of the Channeled Scabland is the Columbia River Basalt Group (CRBG). This continental flood basalt province formed between 17 and 6 Ma, with the majority of the lavas erupting in less than one million years around 16 Ma [Tolan et al., 1989]. The oldest hnnaha Basalt filled in much of the pre-existing topography, beginning the formation of the Columbia Plateau. The bulk of the lavas are contained in the subsequently erupted Grande Ronde Basalt, which is dominantly basaltic andesite in composition. The Wanapum Basalt formed the end of the main plateau-building sequence, and the Saddle Mountain Basalt was erupted over the next 8 million years and was largely confined to filling the ancestral Columbia River channel. The total volume of the CRBG (excluding the related Steens Basalt to the south) is -170,000 km^, covering

164,000 km^ to an average depth of over a kilometer [Tolan et al., 1989]. About 300 Figure 1.3. (a) The lateral extent of the entire CRBG (purple) [after Tolan et al., 1989]. (b) Glacial Lake Missoula (blue) and the Channeled Scabland (green) [after Baker, 1978a]. 23 individual lava flows are recognized, with a typical thickness of 30-40 m and volume of a few hundred to >2000 km^ [Tolan et al., 1989].

There has been considerable scientific debate about the nature of these CRBG eruptions. The initial model of Shaw and Swanson [1970], which argued for turbulent floods of lava and eruption durations of days to weeks, stood unchallenged for more than

20 years. However, based on more recent insights gained from eyewitness observations of lava flows in Hawaii and Iceland, Self et al. [1996, 1997] proposed that most CRBG eruptions lasted for about a decade and involved relatively slow flow underneath an insulating lava crust. In the case of the Roza Member of the Wanapum Basalt, a large array of field observations suggests that the individual lava flows were the product of years of inflation [Thordarson and Self, 1998].

The CRBG emplacement was affected by a number of external factors. The lava flows often followed (and filled to overflowing) the channels of the ancestral Columbia and Snake Rivers. This, and the lack of rain shadowing by the then insignificant Cascade

Range, meant that the lava flows advanced across a very wet landscape. In fact, the classic "entablature and colormade" structure of the CRBG is the result of water-cooling of the lava flows [Long and Wood, 1986]. The lava flows were also diverted around the actively growing Yakima Folds. These anticlinal ridges are influenced by regional tectonics, but have remarkable similarities to wrinkle ridges seen on other planets [e.g.,

Watters, 1988].

Following the emplacement of the CRBG, the Columbia Plateau was not subjected to another major geologic event until the Pleistocene glaciations. Vast deposits of loess. 24 locally reaching 75 m thick, were deposited on top of the lavas [Ringe, 1970]. The

Cordilleran ice sheet advanced southward and encroached upon the northern edge of the

Columbia Plateau, producing moraines and other classic glacial and peri-glacial features

[e.g., Stewart, 1913]. Most importantly, the glaciers also dammed the Clark Fork River, forming glacial Lake Missoula. Glacial Lake Missoula extended across what is now western Montana with an area of approximately 8000 km^ and a volume of ~2000 km^

[Pardee, 1910]. This volume is greater than that of the current Lake Erie and Lake

Ontario combined [Mueller and Mueller, 1997].

The ice dam on glacial Lake Missoula periodically failed, triggering catastrophic floods [Bretz, 1930]. These floods scoured through the loess deposits and deep into the underlying basalts, shaping the Channeled Scabland. Unique geomorphologic features, including streamlined islands, boulder bars, giant current ripples, submerged cataracts, and giant potholes were formed by these catastrophic floods [e.g., Bretz, 1923; Bretz et al., 1956; Baker 1973]. Many of these features are so large that they are difficult to comprehend from the ground and are more readily seen fi"om the air or space [Baker and

Nummendal, 1978]. In fact, despite excellent field observations, the early studies of these catastrophic flood features [Bretz, 1923] were met with extreme skepticism.

The magnitude of the floods is best estimated by modeling the flow through specific locations that must have channeled essentially the entire flux of flood waters (the most spectacular of which is the Wallula Gap near Richland, WA). Peak flood discharge rates were estimated to be on the order of 20x10^ m^/s, with flow velocities as high as 30 m/s 25 and water depths of 60-120 m [Baker, 1971]. Individual floods were thought to have lasted for only a matter of a few to several days [Baker, 1973].

Controversy over the exact number of floods has largely revolved around the interpretation of rhythmically bedded flood deposits. The maximum number of floods is obtained by assuming that each bed records a separate flood [Waitt, 1985, 1987], while a much lower number is obtained by recognizing that an individual flood can have multiple pulses [Baker 1973]. There are at least 5 distinct episodes of flooding from 30 ka to 13 ka, each of which could have involved multiple floods [Baker and Nummendal, 1978].

1.4: Overview of Athabasca Valles, Mars

Like the Channeled Scabland, the martian channel system Athabasca Valles was carved by catastrophic floods [Burr et al., 2002a,b; Berman and Hartman 2002; Plescia

2003] (Fig. 1.4). Located in the Elysium Planitia region of Mars (at ~10°N, 155°E),

Athabasca Valles was first identified by Tanaka and Scott [1986] and Edgett and Rice

[1995] using Viking imagery. There are several different age estimates for the channel system, all derived from impact craters. The impact crater Zunil produced numerous secondaries that litter the floor of Athabasca Valles [McEwen et al., in revision]. If the hypothesis that Zunil is the source of the Shergottite meteorites is correct, Athabasca

Valles must be older than ~2 Ma [McEwen et al., in revision]. Using larger craters, the upper age limit for the surrounding plains is -100 Ma [Plescia, 1990; Lanagan, 2004;

McEwen et al., in revision]. Despite uncertainties in the age of Athabasca Valles, all studies agree that it was carved during the most recent period of Mars' geologic history. 26

Figure 1.4. Mars Orbital Laser Altimeter (MOLA) topography over Athabasca Valles centered at 10°N, 155.5°E [Figure 2 from Burr et al., 2002a]. Contour lines are spaced 0.05 km. The arrows show inferred water flow directions. 27 unlike many of the catastrophic outflow channels that are billions of years old [Baker,

1982].

Another way in which Athabasca Valles is similar to the Channeled Scab land is that it also is thought to have been carved into stacks of mafic flood lavas [Keszthelyi et al.,

2000, in press; Lanagan, 2004] (Fig. 1.5). Some of these lava flows appear to extend at least 1500 km, starting fi-om the Cerberus Fossae (a fissure system), stretching across the

Cerberus plains, continuing through Marte Vallis and entering Amazonis Planitia

[Lanagan, 2004]. Modeling of the emplacement of such immense sheets of lava suggests that it must have had a mafic rheology and that eruption rates were probably an order of magnitude larger than those of typical Columbia River flood basalt flows [Keszthelyi et al., 2000],

Like the lava flows, Athabasca Valles also begins at the Cerberus Fossae [Edgett and

Rice, 1995; Burr et al., 2002a,b; Berman and Hartman, 2002; Plescia, 2003]. This suggests that both water and lava were erupted fi^om the same fissures, but it is not certain that they rose through the same fissure segment, much less simultaneously [Burr et al.,

2002b]. hi any case, water exited the Cerberus Fossae to both sides of the fissures [Burr et al., 2002a; Manga, 2004]. The water then flowed down the topographic gradient to the southwest and carved Athabasca Valles. It was confined on its southeastern margin by a wrinkle ridge, focusing flow into a channel -30 km wide. The flood waters appear to have downcut approximately 70 m into the plains, leaving a series of streamlined islands.

In the lee of some of these islands there are rhythmically bedded layers and dunes [Burr et al., 2002b, 2004]. There are a number of narrow channels carved into the wrinkle Crater MirterMi: Rims and interiors of craters wcll-rcsolvcd in the THEMIS infrared data. Interpreted to have formed by impact events. • Crater Ejecta: Lobate or diffuse deposits surrounding craters.

l^bate Plains: Gently dipping plains wtih distinct legate flow scaips. Interpreted to be lava flows, dominantly from the Elysium rise. The morphology of the lavas are generally consistent with aa-like lava flows. naty-ridged Plains: Rat plains with smooth or platy-ridged surfaces. a I* Interpreted to he sheet-like flffod lava flows of the Cerberus plains. Lobate Scarp: Low topograf^ic scarps with lobate or crenulated traces. Interpreted to be margins of lava flows.

"Clean" Surface: Regicms that have unusually high thennai contrast on small spatial scales. Interpreted to be regions with unusually thin dust cover. The dust cover may be thin because the surface is young (not much time for dust accumulation) or because the dust was recently removed (possibly by water). Channds: Topograhic channels. Interpreted to be formed by aqueous erosion, but some in the lobate plains may be primary volcanic Matures.

Fissures: Topographic fissures. Interpreted to be tensile fractures of the Cerberus Fossae. They may he primary tectonic features or possibly deformation s caused by intrusion of dikes.

Wrinkle Ridges: Line marks trace of steepest part of the ridge. Interpreted to s be folds over thrust faults. Undifferentiated Plains: Smooth or gently undulating surfaces without clear •volcanic or fluvial features. Interpreted to be moderately tectonically deformed lava plains materials. Cenemlly corresponds to wrinkle ridge m

Mantling Deposits: Highstanding surfaces with distinct colian erosion features (especially yardangs). Generally corresponds to the Medusae Fossae Formation. Interpreted to be eroding ash atullor dust deposits.

Knobby Terrain: Isolated high topography, often with small debris aprons. n Interpreted to be remnants of older cratered plains. Figure 1.5. Geomorphic and thermophysical map of the Athabasca Valles region of Mars based on Thermal Emission Imaging System (THEMIS) data [from Keszthelyi et al., 2004a]. K) 00 29 ridge where the flood waters were able to overtop low points [Burr et al., 2002a]. The main channel of Athabasca Valles loses its distinct morphology near the terminus of the wrinkle ridge [Burr et al., 2002b]. Modeling of the flow though Athabasca Valles suggests peak flow rates of order 10® m^/s [Burr et al., 2002a,b, 2004].

The relative timing of aqueous flooding and lava flow emplacement at Athabasca

Valles is difficult to decipher. The floor of the channel preserves both linear ridges, interpreted to be the product of catastrophic flood erosion [Burr et al., 2002a,b], and pristine lava flow textures [Keszthelyi et al., 2004a]. The latter strongly suggest that at least one lava flow moved through the valley after the main charmel-carving flood.

Rootless cones found on the surface of this flow indicate that there could not have been a geologically long period of time between the emplacement of this lava flow and the ground having been soaked by water [Lanagan et al., 2001]. However, as described very briefly in Chapter 4, there are hints of local fluvial or diluvial deposits superposed on the youngest lavas. This may indicate that smaller-scale flooding occurred after the emplacement of the lava flow. It is even plausible that the lava flow was emplaced into the waning stages of a single catastrophic flood.

The observations of Athabasca Valles and its surroundings are made almost exclusively from spacecraft orbiting Mars. The one exception is the radar data that was obtained from Earth-based radio astronomy observatories. The primary data used in this study are Narrow Angle Mars Orbital Camera (MOC) images. MOC was originally flown (as the Mars Observer Camera) on the ill-fated Mars Observer spacecraft [Malin et al., 1992]. Its re-incamation aboard the Mars Global Surveyor (MGS) spacecraft reached 30

Mars in 1997 [Malin and Edgett, 2001]. MOC has a 2048 pixel-wide push-broom imaging system that produces long nadir-looking strips as the MGS spacecraft moves along its polar orbit. While MOC has a wide-angle camera with two color filters, the narrow-angle camera images are all panchromatic. The nominal resolution of the camera is -1.5 m/pixel, but most images are binned in order to improve the signal to noise ratio.

Typical images of Athabasca Valles are 4-6 m/pixel. However, recently the MGS spacecraft has experimented with both roll and pitch maneuvers. Rolled images allow for improved stereo observations and pitching while imaging allows "super-resolution" in the down track direction. A 0.5x1.5 m/pixel image over Athabasca Valles was obtained using this technique. 31

CHAPTER 2: OROGENIC TECTONISM ON lO

This chapter is a modified version of the published paper "Orogenic tectonism on lo" by W. L. Jaeger, E. P., Turtle, L. P. Keszthelyi, J. Radebaugh, A. S., McEwen and R. T.

Pappalardo [Jaeger et al., 2003]. As such, it is the result of the combined efforts of multiple researchers, and is only one piece of a larger set of collaborative publications.

In an effort to identify my original work and to acknowledge my coauthors for their input, the contributions of each author are specified here. Elizabeth Turtle and Alfi-ed

McEwen provided funding for this research. Additionally, Elizabeth Turtle contributed to the hotspot swell model and to the overall fabric of this paper, and Alfired McEwen has helped to put this work into the larger context of previous and ongoing lo research.

Laszlo Keszthelyi reviewed the statistical analysis in section 2.2 and wrote a FORTRAN code to solve Equation 2. He has also focused on the response of volcanism to tectonics on lo [Keszthelyi et al., 2004b]. Jani Radebaugh cataloged ionian patera and combined that data with the mountain list in this publication (Table 2.1) to make a larger database

[Radebaugh et al., 2004]. She also provided an independent check of associated mountains and paterae. Bob Pappalardo reviewed the stress calculations in section 2.3.1 and recommended the use of Byerlee's Law in Figure 2.7. I did the remaining work in this chapter, which includes, but is not limited to, cataloging mountains (and associated paterae), estimating the total number of ionian mountains, determining the probability of the observed mountain-patera associations, reviewing previous mountain formation models, calculating the average stress state of lo's lithosphere, computing the volumetric 32 strain resulting from subsidence and thermal stresses, estimating the volume of material available to build mountains, placing a lower limit on lo's lithospheric thickness, conceptualizing the hotspot swell model, and interpreting structural features on and around mountains.

2.1: Introduction

The surface of lo, Jupiter's iimermost Galilean satellite, is comprised of lofty mountains and low-relief volcanic centers separated by extensive plains. The mountains are observed to be rugged, isolated peaks that are approximately randomly distributed across the surface [Carr et al., 1998], though subtle concentrations of mountains may exist near 25°N, 65°W and 20°S, 265°W [Schenk et al., 2001]. Stereoscopic and shadow measurements of these peaks yield heights ranging from a few kilometers to approximately 18 kilometers [Schenk et al, 2001]. Because of their soaring heights, the mountains are inferred to be dominantly silicate structures rather than sulfiar-rich protuberances as such volatile compounds should not support steep topography much in excess of 1 kilometer [Clow and Carr, 1980]. hnages returned by the Voyager spacecraft in 1979 revealed that caldera-like depressions commonly abut mountains [Masursky et al., 1979]. This finding initially led researchers to speculate that lo's mountains were volcanic in origin [Masursky et al., 1979; Carr et al., 1979]. However, subsequent studies have made convincing argiunents that, while a few peaks may be volcanoes [e.g., Moore et al., 1986], the mountains are predominantly tectonic massifs [Schaber, 1982; Nash et al., 1986; McEwen et al., 1989]. 33

Images taken by Voyager and by the Solid State Imaging (SSI) instrument onboard the Galileo spacecraft have greatly improved our knowledge about mountains on lo; however, their mechanism of formation is not yet well understood. Schenk and Bulmer

(1998) suggested that the rapid volcanic resurfacing of lo [e.g., Johnson et al., 1979] generates a horizontal compressive stress in the lithosphere that is sufficient to uplift the ionian moimtains via thrust faulting. Alternatively, McKinnon et al. [2001] proposed that the lithospheric heating associated with a sustained reduction in lo's volcanic activity on a local, regional or global scale could induce a large compressive stress at the base of the lithosphere such that over time, as resurfacing rates wax and wane, the fluctuating thermally-induced stress could lead to alternating episodes of compressive and tensile faulting. They speculated that, as a consequence of repeated normal and reverse faulting, lo's surface might be similar to the chaos terrain of Europa with coherent crustal blocks

(i.e., mountains) "floating" in a matrix of highly disrupted material.

The extent to which lo's mountains are associated with volcanic features is disputed.

Using Voyager and early Galileo data, Carr et al. [1998] found no correlation between mountains and hotspots or plume eruptions. They did note, however, an apparent association between mountains and paterae (irregular or scalloped depressions), which they assumed synonymous with volcanic calderas. A more recent study by Schenk et al.

[2001] investigated the global distribution of mountains and volcanic centers. Their results suggest a hemispheric-scale anticorrelation between the densest concentrations of these two feature types. McEwen et al. [2000] examined images taken by the Galileo spacecraft during its 124, 125, and 127 flybys. They found that, of the 13 mountains 34 imaged at resolutions higher than 0.5 km/pixel, 6 are incised by paterae. These mountains are located within a region of relatively high mountain density and low volcanic center density [Schenk et al., 2001].

hi this chapter, 100 well-imaged ionian mountains are cataloged in order to quantify their relationship with paterae. The fraction of those mountains that have characteristics suggesting a primarily tectonic origin as well as the fraction that appear to be primarily volcanic constructs, are determined based on specific criteria. Next, the relative importance of subsidence and thermal stresses on ionian mountain formation is computed as a function of lithospheric thickness; using these results, a lower limit is placed on the thickness of the thermal Uthosphere. Finally, a mechanism for mountain formation on lo in which rising asthenospheric diapirs dome the overlying Uthosphere and focus the horizontal compressive stresses is proposed. This model is then shown to be consistent with the isolated occurrence of mountains and their association with paterae.

2.2: A Link Between Tectonism and Volcanism on lo

As mentioned above, several studies have hinted at a juxtaposition of lo's mountains and paterae (which commonly have been referred to as calderas) [Masursky et al., 1979;

Carr et al., 1998; McEwen et al., 2000]. However, Schenk et al. [2001] found an anticorrelation between the global distributions of mountains and volcanic centers

(mostly paterae). These apparently contradictory results may, in fact, be consistent with one another; mountains and paterae may be associated on a local scale but not on a global scale. 35

Before presenting statistics on mountain-patera associations, it is necessary to specify how the term "patera" is used in this manuscript. The hitemational Astronomical Union

(lAU) defines a patera as "an irregular crater, or a complex one with scalloped edges."

Recurrently in lo literature the term caldera has been used interchangeably with patera in referring to a topographic depression >1 km in diameter that is not obviously impact- generated. Despite this common usage, the term caldera implies a mode of formation that may not apply to paterae [Radebaugh et al., 2001]. Paraphrasing the definition Wood

[1984] suggested for planetary application, a caldera is a multikilometer-wide, quasi- circular depression that is the surface expression of a roof collapse over a partially evacuated, shallow magma chamber. Calderas are primarily volcanic in origin, and their formation is typically syn- and immediately posteruptive. Lipman [2000] describes similar featiires known as "volcano-tectonic depressions" (VTDs) as topographic lows with variable subsidence geometries formed when eruptions vent fi-om regional graben.

VTDs are formed by a combination of volcanism and regional tectonism. On lo, we do not yet have the evidence to definitively classify paterae as calderas, since we cannot ascertain the presence of shallow, precollapse magma chambers under paterae. However, evidence suggests that some ionian paterae may be VTDs. In several instances the locations of paterae and/or their perimeter morphologies are tectonically controlled

[Radebaugh et al., 2001]. Because there are numerous hypotheses and perhaps multiple mechanisms for the formation of lo's caldera-like depressions, the non-genetic term

"patera" is used throughout this discussion. Ionian paterae (Fig. 2.1) can appear either as 36

Figure 2.1. A small sampling of Ionian paterae intended to illustrate the varied nature of these landforms. Scale bars are 50 km. (a) Chaac Patera (11.8°N, 157.2°W) is a green- floored topographic depression with irregular margins that is nested within a larger, roughly elliptical light-floored patera. The smaller, white-floored feature to the east is Balder Patera, (b) Hi'iaka Patera (3.rS, 79.8°W) is a depression with relatively straight margins. The eastern half of its floor is dark, indicating recent silicate lava flows. There is a light-floored, butterfly-shaped patera in the lower left comer of the image, (c) Tupan Patera (19°S, 14rW) has a mottled floor with darker lava flows apparently surrounding a central island, (d) Emakong Patera (3°S, 120°W) is surromded by radially oriented lava flows. The lava appears to have flowed away from the patera; therefore, the margin of the patera is inferred to be a local topographic high. 37 dark-floored depressions (indicating the presence of young silicate lava flows) or as light- floored depressions that are coated by sulfurous volatiles [Radebaugh et al., 2001].

Table 2.1 lists the 151 ionian mountains reported by Carr et al. [1998], Schenk et al.

[2001], and this study, noting their general morphology and proximity to paterae. The positions of adjacent paterae are also noted. In cases where mountains reported by Carr et al. [1998] and/or Schenk et al. [2001] are not apparently included in the table, they may be (1) listed under slightly different coordinates in this study, (2) features that are suspected to have been listed under two or more sets of coordinates in the original paper and are cataloged only once here, or (3) individual peaks within a larger massif where only the larger feature is recorded here. In this context, the term "mountain" is used to indicate any discrete feature that can be determined by its shadow or in limb images to have an elevation significantly (>1 km) above that of the surrounding terrain; thus it includes features classified according to lAU nomenclature as montes, mensae, tholi, and plana. In some instances paterae surrounded by radial lava flows were inferred to be local topographic rises, but these are not included in Table 2.1 because their slopes are too shallow to cast shadows even at large solar incidence angles. Of the 151 reported ionian mountains, 145 are verified by this study. Furthermore, 101 of the 145 verified mountains reside in regions of lo that were imaged by Voyager and/or Galileo at moderate or high resolutions and at relatively large incidence angles, which emphasize topography (Fig. 2.2). In these regions, which comprise 58% of lo's surface area, all existing mountains are likely to have been identified. Extrapolating from this area, the 38

TABLE 2.L Comprehensive list of ionian mountains, mesas, plateaus, small peaks and volcanic constructs.

Mountain Position Identified by: 0 Patera(e'> Positionfs) 1 O "I W) 00 rsi •S I W) dj o^ON cd Q 0 w *3 (D Q o 1 ao y oji: O d Feature name hJ u 00 ;z Lat. (deg), Lon. (deg) ^38.2 "2.8 • • T ^30.4 "7.3 • • T 0 1^86.0 "8.0 • • ^36.6 "10.1 • • • T 1 34.9, 11.9 -12.5 14.7 • • T 0 ^19.9 "24.2 • • 0 yj ^31.2 "24.7 • • • 1^-24.0 "25.0 • • • T 1 -21.8, 24.3 P-35.4 "25.7 0 • T -84.9 28.9 • T P-10.4 29.1 0 • T 1 -13.5,23.9 Ethiopia Planum -44.1 30.0 • • • T 0 P35.O "30.1 • • 0 rj ^30.1 "31.8 • • • T 1 (or 2) 29.4, 22.3 (31.5, 34.6) -27.2 33.8 • • • T Pan Mensa -50.2 34.4 • • • T 1 (or 2) -47.5, 37.8 (-51.9, 32.2) P7O.O "36.0 • • T P-87.4 "36.4 • • • T "38.7 "40.5 • • T 0 "18.3 "43.5 • • • T 0 "-25.4 "43.7 • • • T "-65.0 "43.8 • • T 0 "12.8 "46.2 • • • T Haemus Montes -69.7 47.7 • • • T 1 -68.5, 58.2 "69.9 "50.6 • • T "-11.9 "55.8 • • T 0 "-38.9 "56.2 0 YT "18.9 "56.6 • • T 0 39

TABLE 2.1 - continued

Mountain Position Identified bv: PateraCe) Position(s) I Ba O S.s o g .a CiX) ooON

TABLE 2.1 - continued

Mountain Position Identified bv: o Paterafe) Position(s) Ba o o CJ) 0\00 fS •S I 60 a\ U o w cd ^ 6 cd -M o 3 •«-> 3 'S <« B 2 o •c o y O o0) 6 Feature name hJ J u CO H :z Lat. (deg), Lon. (deg) Tvashtar Mensae ^58.8 122.4 • • T 2 61.5, 120.2; 64.8, 126.0 -13.0 124.2 • T 0 -26.3 124.5 • • T 0 45.6 125.9 • • T 1 47.9, 123.5 Euxine Mons 26.5 126.2 • • T 2 25.6, 124.3; 23.3, 125.1 P14.5 127.2 • • T 1 16.3, 126.1 M2.5 P129.7 • • T 1 -41.6, 137.5 Seth Mons -10.3 134.0 • • T 0 P-19.4 ^148.6 0 T 1 -18.7, 150.7 P-37.0 ^148.7 0 T 1 -35.0, 152.0 22.4 151.6 • T 1 21.1, 151.6 P-64.2 ^157.2 0 VT ^0 20.3 157.3 T 1 19.4, 158.8 28.9 159.5 T 2 27.5, 157.9; 31.6, 159.1 P-68.2 ^159.5 0 T 0 Tohil Mons -28.9 160.3 • • T 3 -35, 152; -27, 158;-28, 160 13.4 160.7 T 2 11.3; 155.8; 11.8, 157.2 P-60.8 "161.6 0 T 47.0 162.0 • • T 0 25.2 165.5 V 1 25.2, 165.5 -23.9 171.9 • • T 0 16.8 173.7 V 1 16.8, 173.7 18.7 174.4 V 0 ^64.5 "174.8 • • T -77.7 179.0 0 T -1.7 183.7 T 0 4.6 185.8 T 1 6.3, 187.6 20.7 188.7 • • T 0 Dorian Montes -21.5 193.0 • • T 2 -23.0, 193.9; -20, 194 59.2 195.5 • • T 1 59.4, 198.6 -73.1 196.8 • • T "73.8 "200.5 • T 41

TABLE 2.1 - continued

Mountain Position Identified bv: o PateraCe) Position(s1 I ac o o o 00 •S I (DCtO a\

TABLE 2.1 - continued

Mountain Position Identified bv: u Patera(e") Position's) Ba IO o .a o &0 00 Tholus -10.9 348.7 • V 1 -10.9, 348.7 Inachus Tholus -15.8 348.8 • V 1 -15.8, 348.8 Echo Mensa -79.9 355.7 c • T 0 Sources of data; " Carr et al. [1998]; ^ Schenk et al. [2001]. Y = categorization is inferred from, but not explicitly stated in, Schenk et al. [2001]. • = positively identified mountain; o = tentatively identified mountain. Note; Values for longitude increase to the west. 43

360 300 240 180 120 60 0

Longitude (degrees)

Figure 2.2. A plot of all ionian mountains reported by Carr et al. [1998], Schenk et al. [2001], and this study [c.f. Table 2.1]. Verified mountains are plotted as •s, while features reported, but not verified are plotted as Xs. The shaded area indicates the 58% of lo that is imaged at high enough resolutions and large enough solar incidence angles for topography to be easily seen. 44 total number of mountains on lo's surface is estimated to be about 174, suggesting that less than 30 mountains have yet to be identified.

The 101 mountains that reside in well-imaged areas were classified as either

"associated" or "not associated" with one or more paterae and as either "volcanic" or

"tectonic" in origin. In this study, a mountain and patera are "associated" only if they are in direct contact (Fig. 2.3). Mountains that are approximately axially symmetric with conic to shield-like morphologies and summit depressions or lava flows issuing from their summits were classified as volcanic constructs (Fig. 2.4), and rugged, asymmetric mountains were classified as tectonic massifs (Fig. 2.5). Table 2.2 and Figure 2.6 show that only 4 of the 101 mountains are apparent volcanic constructs while 97 are either tectonic in origin or have been tectonically modified subsequent to their formation.

Additionally, at least 40 of the 97 tectonic massifs (or >41%) are associated with paterae; the presence or absence of paterae was not determined for 6 mountains in the well- imaged region (Fig. 2.6). Schenk et al. [2001] tabulate similar statistics for the number of mountains in close proximity to paterae (45%). Either percentage is too high to be coincidental. In a complementary statistical investigation, Radebaugh et al. [2001] find that only 13% of paterae have mountains associated with them. Thus paterae can readily occur independently of mountains, but it is less common for mountains to occur independently of paterae.

On average, a patera and a mountain will be in contact if their geometric centers are closer together than the sum of their average radii. The fraction of lo's surface area in 50 km

Figure 2.3. A mountain in contact with Shamshu Patera (9.8°S, 63.6°W). 46

Figure 2.4. Examples of mountains classified as volcanoes, (a) Two shield volcanoes located near the plume Zamama at 16.8°N, 173.7°W and 18.7°N, 174.4°W. (b) Apis Tholus (10.9°S, 348.7°W) and Inachus Tholus (15.8°S, 348.8°W) are substantially larger than other ionian volcanoes. 47

50 km

:P

mWi; ^ m,4i

k

Figure 2.5. Examples of mountains classified as tectonic massifs, (a) The Hi'iaka Montes (7.4°S, 78.7°W and 2.rS, 82.3°W) are elongate features with relatively flat, ridged surfaces. The exceptions to this generalization are the northern peak, which is quite rugged, and the southernmost margin, which appears to have been modified by a landslide, (b) Shamshu Mons (11.3°S, 71.7W) is an asymmetric mountain that consists of two parts, a large, hummocky mass to the southeast and a narrow ridge to the northwest, separated by a valley. 48

TABLE 2.2. Characteristics of the 101 moimtains investigated.

Description Number

Mountains included in statistical analysis (i.e. located in well-imaged regions) 101

Volcanoes 4

Tectonic massifs 97

Tectonic massifs "associated" with paterae 40

Tectonic massifs "not associated" with paterae 51

Tectonic massifs with uncertain relationship to paterae 6

Paterae residing in regions well-imaged for mountain detection ^296

Paterae touching tectonic mountains 52-53 ''From Radebaugh et al. [2001], 49

360 300 240 180 120 60 0

Longitude (degrees)

Figure 2.6. The locations of all Ionian mountains reported by Carr et al. [1998], Schenk et al. [2001], and this study [c.f. Table 2.1]. Mountain subtypes are symbolically represented in the following way: volcanoes (A), tectonic mountains that contact paterae (•), tectonic mountains that do not contact paterae (•), and mountains whose relationship to paterae was not determined (?). The gray shading highlights the 58% of lo that was well-imaged for mountain detection. 50 which the center of a mountain can be located such that any part of that mountain will be in contact with a patera (/) can be estimated by

AnR' ' ^ ^ where rip is the number of paterae on lo, is the average basal radius of a mountain, Vp is the average radius of a patera and R is the radius of lo. This equation makes the simplifying assumptions that mountains have circular bases and that paterae do not overlap; nevertheless, it gives a reasonable approximation. If only the 58% of lo for which there is sufficient coverage to identify all mountains is considered, the denominator in Equation [1] becomes 2.32 In this region Radebaugh et al. [2001] find the number of paterae, Up, to be 296, and they determine the average patera area to be 0 9 2119 km . This corresponds to a patera radius, rp, of 26 km. Assuming 12,080 km to be the average basal area of a mountain [Schenk et al., 2001], the average mountain radius, rm, is 62 km. For these values, / is 0.298. Next, the probability (p) that x out of rim randomly placed mountains will be located within / (i.e., will contact paterae) is determined using the following equation:

yt I p't. ,, '' ./[rxc-/)"-'"] p-21

[Barr and Xehna, 1983]. By this calculation, the probability that >40 of the 97 tectonic mountains will contact paterae, as is observed, is <1.08%. Taking into account the fact that some of lo's paterae are nested (e.g. Tvashtar Catena, Chaac Patera) will further reduce the probability. Equations [2.1] and [2.2] do not address the fact that several of 51 the tectonic mountains (about 30% of those touching paterae) abut more than one patera.

The spatial association between mountains and paterae on lo suggests a genetic link between the two features (and, therefore, between tectonism and volcanism). Faults associated with mountain uplift may act as conduits for magma ascent, thus facilitating patera formation in the vicinity of mountains. This hypothesis is revisited after examining the details of mountain formation.

2.3: Orogenic Tectonism on lo

The process by which ionian mountains form is not well understood. The rapid resurfacing rate of lo (~1 cm/yr on global average [Johnson et al., 1979; Blaney et al.,

1995; Phillips, 2000]) may be partly to blame for the uncertainty. Lava flows and sulfiirous plume deposits quickly blanket the moon's surface, obscuring tectonic features.

For this reason, researchers are forced to rely heavily on a theoretical approach when attempting to understand lo's orogenic cycle. The following section reviews and ftirther analyzes previous mountain formation models, draws new inferences about the thickness of lo's thermal lithosphere, and presents a possible mechanism for focusing lithospheric stresses to uplift isolated mountains.

2.3.1: Analysis of Previous Models

Schenk and Bulmer [1998] reason that the rapid volcanic resurfacing of lo must generate a horizontal compressive stress in the lithosphere as concentric shells of material 52 are continually buried, decreasing their effective radii. At depth z, this compressive stress

{(^subsidence) IS given by,

^subsidence 1-vi ^ Rn ' where E is Young's modulus (~10'' Pa), vis Poisson's ratio (~0.25) and R is the radius of lo (1821 km, on average) [McKinnon et al., 2001]. This does not account for closing of pore spaces with increasing pressure [Leone and Wilson, 2001]. Schenk and Bulmer

[1998] propose that this subsidence-induced stress leads to thrust faulting that could uplift mountains. As supporting evidence for this hypothesis, they show that the morphology of East Euboea Mons (48°S; 336°W) is similar to that of terrestrial basement-cored thrust sheets. Using equation [2.3] and assuming that the compressive strength of lo's brittle lithosphere at the surface is <275 MPa (the maximum value for the compressive strength of unconfined rock samples of gabbro, diabase, basalt, amphibolite and peridotite

[Ahrens, 1995]), under globally uniform resurfacing CTsubsidence is large enough to cause brittle failure at a depth of ~4 km (Fig. 2.7). This finding has two important implications;

(1) it indicates that lo's lithosphere is pervasively fi*actured because the available stresses are larger than the strength of the rock for all but the upper few kilometers, and (2) it implies that the compressive stress in lo's lithosphere is governed by fnctional sliding below ~4 km.

Turtle et al. [2001] use two-dimensional finite element modeling to predict the distribution of mountains formed by the Schenk and Bulmer [1998] model. They find 53

0

Coulomb failure at ~4 km 5

10

15

20

25

30 0 1000 3000 4000 5000 Stress (MPa)

Figure 2.7. The compressive strength of the ionian Uthosphere and the magnitude of the compressive horizontal stress {(Tsubsidence + CTthermai) as a function of depth for a 30 km- thick Uthosphere. The maximum compressive strength of lo's Uthosphere at the surface is assumed to be 275 MPa (the maximum compressive strength of unconfined rock samples of gabbro and several other rock types [Ahrens, 1995]), thrust faults are assumed to form at 30°, and the resurfacing rate is assumed to be 1 cm/yr. The resulting equation for Coulomb failure is r = 79.4 + 0.58cr„, where r is the critical shear stress (in MPa) and cr„ is the normal stress (in MPa). On lo, the compressive horizontal stress due to subsidence is sufficient to initiate faulting at a depth of ~4 km. Once faults exist, surplus stress should be relieved by slip along these preexisting planes of weakness. According to Byerlee's Law, fiictional sliding is governed by the equation = 0.85cr„, where Zfis the shear stress at which sliding begins. The solid black line illustrates the approximate stress state of the Uthosphere and the shaded regions show the surplus stress (i.e., that which is available to drive mountain uplift). The light gray region shows the magnitude of the compressive stress due to subsidence and the dark gray region shows the contribution of thermal expansion. 54 that mountains formed by thrust faulting in a pervasively fractured lithosphere will be uplifted as parallel mountain ranges rather than as the isolated peaks that are observed.

However, they note that taking into account factors such as non-uniform resurfacing and lithospheric heterogeneities could alter their findings.

McKinnon et al. [2001] propose an alternative model for mountain formation on lo.

They suggest that a decrease in the rate of volcanic resurfacing will result in heat build­ up at the base of the lithosphere. This can occur on a local, regional or global scale. The rise in temperature of the lower lithosphere generates a horizontal compressive stress

{(Tthermai) that is given by

EaAT ^thermal ~ > ' L^-^i l-V where E is Young's modulus, a is the linear coefficient of thermal expansion, AT is the increase in temperature of the lowermost lithosphere and vis Poisson's ratio [McKinnon et al., 2001]. They propose that the compressive stress in the lower lithosphere is taken up along thrust faults that often penetrate into the upper lithosphere. Displacement along these faults can exceed what is needed to decompress the upper lithosphere and, as a result, induce a tensile stress. Their calculations indicate that the magnitude of this tensile stress is sufficient to cause normal faulting of the upper lithosphere along new and reactivated fault planes. Continued heating thins the lithosphere and the resulting thermal expansion acts to uplift and tilt the previously rotated lithospheric blocks. Thus, according to their model, the surface of lo consists of isolated, coherent lithospheric blocks in a matrix of "chaotic terrain." 55

The model presented in McKinnon et al. [2001] calls attention to several key issues.

The volcanic resurfacing rate on lo should indeed fluctuate laterally and temporally, although it also may be quite uniform when averaged over relevant timescales [Carr et al., 1998]. Heating and thermal expansion will generate large stresses at the base of the ionian lithosphere irrespective of fluctuations in the resurfacing rate. These thermal stresses are the result of material originally on the surface of lo at -100 K [e.g.. Spencer and Schneider, 1996] being buried, via continued resurfacing, to the base of the lithosphere where we estimate the temperature to be about 1500 K. In terrestrial literature the base of the thermal lithosphere (i.e., the interface between conduction and convection) is often taken to be the 1600 K isotherm [e.g., Turcotte and Schubert, 1982], but under the lower gravity (pressure) of lo, lithospheric rocks would be largely molten at

1600 K [Keszthelyi et al., 1999]. Keszthelyi and McEwen [1997] predict that 10% partial melting of a relatively undifferentiated lo is achieved at a temperature of -1440 K.

So a temperature of-1500 ± 100 K seems a reasonable estimate for the isotherm defining the base of the ionian thermal lithosphere. The relatively small fluctuations in resurfacing rate discussed in McKinnon et al. [2001] will produce perturbations on this already existing, thermally induced stress.

In this study the magnitude of the thermal stress that results from the steady-state subsidence of lithospheric material was computed using Equation [2.4] and assuming the lithospheric temperature profile of O'Reilly and Davies [1981]. For an average resurfacing rate of -1 cm/yr [e.g., Johnson et al, 1979], an increase in temperature of

1400 K will result in a thermal stress on the order of 2 GPa. This corresponds to -3% 56 thermal expansion of the rock. Perturbations to the resurfacing rate that are sustained for about ten thousand years can lead to some variation in this value, but it would take a major decrease in resurfacing sustained for nearly one hundred thousand years before heat conducting from the base of the lithosphere could significantly alter this average value.

Local perturbations of the lithospheric temperature profile must exist in close proximity to magma conduits. However, one of the major findings of the Galileo

Mission has been that the locations of volcanic vents are remarkably fixed in space over timescales of decades [McEwen et al., 1998]. Furthermore, the low-albedo areas thought to be associated with volcanic vents make up a small fraction (<1%) of lo's surface

[McEwen et al., 1985] implying that the volcanic plumbing makes up a similarly small volume fraction of the lithosphere. These observations suggest that magma conduits should perturb the temperature profile for a few percent of lo's lithosphere. This is a relatively small effect compared to the heating at the base of the lithosphere, but it does mean that thermally induced stresses in the upper part of the lithosphere are slightly underestimated.

Ignoring the localized effects of intrusions and assuming a resurfacing rate of 1 cm/yr, the peak stress generated by thermal expansion will exceed that caused by subsidence for lithospheres thinner than about 40 km. However, the integrated effect of thermal expansion is not as great as that of subsidence. The subsidence-induced stress is large for much of the lithosphere as Figure 7 illustrates while, for any of the estimated resurfacing rates [Johnson et al., 1979; Carr, 1986; Blaney et al., 1995; Phillips, 2000], 57 only the lowermost lithosphere is significantly heated [O'Reilly and Davies, 1981],

Thus, while the stresses produced by thermal expansion can be locally larger than those produced by subsidence, they act on a much smaller fraction of the lithosphere. The region with the most thermal expansion is the partially molten upper asthenosphere where the temperature is in excess of ~1500 K. However, in this region the rock loses its mechanical strength and behaves as a viscous fluid on the time scale of interest. Thermal expansion of the plastic asthenosphere will produce a very small amount of extension of the overlying material rather than contributing to the compressive deformation of the lithosphere. In the following subsection, the cumulative effects of both thermal and subsidence stresses on the ionian lithosphere are quantified and new constraints are placed on the thickness of lo's lithosphere.

2.3.2: Estimating the Thickness oflo 's Lithosphere

While not obvious from equations [3] and [4], the effect that subsidence and thermal stresses have on lithospheric deformation is strongly influenced by the thickness of the lithosphere. hi this section, the volume of material that should be vertically displaced by each process is calculated as a function of lithospheric thickness (i.e., areal strain is integrated with depth). By comparing the total volume of lo's tectonic mountains with the volume of material uplifted by subsidence and thermal stresses, a minimum lithospheric thickness is determined.

The change in volume caused by subsidence of material from the surface of lo to the base of the lithosphere {AVsubsidence is given by 58

subsidence = ^ d - \ A TJZ^ dz, [2.5] R-d where R is the radius of lo, d is the thickness of lo's Uthosphere, and z is depth. The fact that lo is a tri-axial ellipsoid, rather than a perfect sphere, introduces a negligible correction factor that is ignored in these calculations. The thermal expansion of a thin spherical shell of material is found by

V ={\ + aMf -Anr^ dr, [2.6] where V is the new volume, a is the linear coefficient of thermal expansion, AT is the change in temperature, r is radius and dr is the thickness of the shell. Thus, the thermally induced change in volume of the entire Uthosphere (AVthermai) is given by

4z(R-zyx {2a T(z)+ [a r(z)f }dz, [2.7] 0 where T is temperature. For a rapidly resurfacing Uthosphere with conductive heating only at its base, O'Reilly and Davies [1981] show that

e -1 where To is the surface temperature of lo, Tg is the temperature at the base of the thermal

Uthosphere (i.e. the boundary at which the timescale for subsidence equals the timescale for convection), and I is the ratio of thermal diffusivity to subsidence rate. Part of this volume change is accommodated by compression of the lithospheric rocks. This volume change (AVcompression) is given by 59

IV ^Kompression=- j ^ 7iz' X {s^ + £^ + S,) dz, [2.9] R-d where Sx, Sy, and ^ are the values for strain in the x, y and z directions, respectively.

Turcotte and Schubert [1982] give the following equations for strain in three dimensions:

I V V r„ s^= — (T^ CTy CTj [2.10] 'EE^E V 1 V s = cr + —cr cr^ [2.11] 'EE'E V V 1 Sz= O'x — [2.12] EE^E

In Equations [2.10], [2.11] and [2.12] vis Poisson's ratio, E is Young's modulus and o^,

(7y and (Jz are the stresses in the x, y and z directions, respectively. The horizontal stresses cTx and <7y are equal to the sum of (Jsubsidence, (^thermal (until they exceed the strength of the rock) and the horizontal component of lithostatic loading, cXoverburden, which is given by

_ l-2v ^overburden ^^P S [2-13] 3(l-v)

Here p is density, g is acceleration due to gravity, and z is depth. The vertical stress cTz is lithostatic (i.e. cTz = pgz).

The analytical solutions for Equations [2.5] and [2.7] were computed as was a numerical solution for Equation [2.9]. The results are plotted as a function of lithospheric thickness in Figure 2.8. The base of the thermal lithosphere was taken to be the 1500 K isotherm in these computations. The volume change induced by thermal stresses is shown for three different resurfacing rates; however, the volume accommodated by the 60

Subsidence Thermal - 1 mm/yr Thermal - 1 cm/yr - - -Thermal - 10 cm/yr 10 - Compression

15 -

•§. 30 -•

35 •

40 •

45 -

•1 0 1 2 3 4 5 6 Volume change (x 10 7 km3 )

Figure 2.8. The volume of material displaced by subsidence stresses (blue trend) and thermal expansion (red trends), and the extent to which the compressibility of lithospheric rock (green trend) can counteract those processes are plotted as functions of lithospheric thickness. Positive values for AV indicate that material is displaced by compressive stresses (i.e., uplifted). The following numbers were used in the computation of these trends: the linear coefficient of thermal expansion = 1.5 x 10"^ K"' [Touloukian et al., 1989], the temperature at the base of the lithosphere = 1500 K, the temperature of the surface = 100 K [e.g., Spencer and Schneider, 1996], thermal diffusivity = 4 x 10'^ m^/s. Young's modulus = 1 x 10^^ Pa, and Poisson's ratio = 0.25. The solid red trend was calculated for the generally accepted globally averaged resurfacing rate of 1 cm/yr [e.g., Johnson et al., 1979]. The other two red trends bracket the uncertainty in resurfacing rate with the dashed line corresponding to 10 cm/yr and the dotted line to 1 mm/yr. For a resurfacing rate of 1 cm/yr, subsidence is the dominant factor driving mountain uplift if the lithosphere is thicker than ~12 km. 61 compression of lithospheric rock was computed using a resurfacing rate of 1 cm/yr.

From Figure 2.8 it can be seen that the volume change induced by thermal expansion is dominant only if the lithosphere is thin or if resurfacing rates are low. For thicker lithospheres and/or higher resurfacing rates, subsidence-related compression dominates.

At the generally accepted average resurfacing rate of 1 cm/yr [Johnson et al., 1979; Carr,

1986; Blaney et al., 1995; Phillips, 2000], AVthermai exceeds AVsubsidence only if the lithosphere is thinner than ~12 km. On the other hand, if the resurfacing rate were as low as 1 rmn/yr, AVthermai would exceed AVsubsidence for a lithosphere as thick as ~34 km. lo undoubtedly has heterogeneities in its resurfacing rate, and the possibility that ionijin mountain formation is largely driven by these fluctuations as McKinnon et al. [2001] suggest cannot be ruled out. Using Voyager and Galileo data, Phillips [2000] examined spatial variations in recent resurfacing. She found that the resurfacing rate varies by more than an order of magnitude depending on geologic setting. However, it is conceivable that styles of volcanism and locations of volcanic centers are temporally variable enough to smooth regional fluctuations in the resurfacing rate over time scales relevant to the thermal structure of the lithosphere (-10^ years). Quantifying temporal fluctuations in the globally averaged resxirfacing rate over these time scales is not currently possible, but substantial variations are unlikely because they require the resurfacing rate averaged over ~10^ years to be significantly different from the estimates derived by averaging over ~20 years (from comparison of Voyager and Galileo datasets

[Phillips, 2000]) and -10® years (from the absence of impact craters [Johnson et al.,

1979]). Additionally, the paucity of sizable volcanic constructs on lo is consistent with 62 eraptive centers being migratory or transient on intermediate timescales. These findings suggest that thermal stresses should be significant, but not dominant, under the most likely circumstances.

The total volimie of material that must be pushed up from depth as determined from

Equations [2.5], [2.7], and [2.9] can be compared to the total volume of ionian moimtains.

Schenk et al. [2001] find the average length, width, height, and basal area of a mountain to be about 158 km, 80 km, 6.3 km, and 12,080 km^, respectively. Assuming mountain geometries are intermediate to triangular right prisms and rectangular right pyramids, the average volume of an ionian mountain is -32,500 km^. If there are -165 mountains on lo's surface, then the total volume they vertically displace from the lithosphere is -5.4 x

106 km 3 . This volume corresponds to a lithospheric thickness of-12 km (Fig. 2.9).

This estimate of lithospheric thickness is a minimum for two reasons: (1) compression can be accommodated in ways other than mountain uplift, and (2) the base level from which topography is measured may be too high. Leone and Wilson [2001] suggest that the uppermost layers of lo have -30% porosity, a number that rapidly decreases with depth due to compaction. The decrease in porosity could accommodate a significant volume of material. Similarly, lo's copious volcanism suggests the presence of numerous magma chambers within the crust. Compression of magma chamber void space could also reduce the volume uplifted as mountains. Much volume can also be concealed in broad, regional uplifts with shallow topography. Gaskell et al. [1988] initially reported such basins and swells using Voyager data, but they have not been 63

a 30

Volume change (x 10 7 km3 )

Figure 2.9. Estimated volume of uplifted material as a function of lithospheric thickness. The trend shown in this plot is the sum of the blue, green, and solid red trends in Figure 8. The 165 projected ionian mountains should contain roughly 54 million cubic kilometers of material. This corresponds to a lithospheric thickness of ~12 km, and we consider this value to be a lower limit. confirmed by Galileo data [Thomas et al., 1998]. Continued detailed examination of the shape of lo [e.g., Oberst and Schuster, 2004] may better quantify this factor.

A more intractable uncertainty in these calculations is the selection of a base level.

Topography must be measured with respect to a specific elevation, and in the above computations that base level is taken to be the plains inmiediately surrounding a mountain. However, mass wasting may deposit an apron of material on those plains causing measurements of mountain heights to be too low and, in turn, mountain volumes and lithospheric thickness to be underestimated. Likewise, these calculations do not take into account the possibility that there is a patchy, perhaps kilometer-thick, mobile layer of sulfixrous volatiles that does not subside. If this is the case, the level from which topography should be measured (and at which basal areas of mountains should be measured) is the elevation at the base of this volatile layer.

While 12 km should be viewed only as a lower limit, these computations can also place a weak upper limit on lithospheric thickness. For example, nearly 48 million km^ of material (the equivalent of a hemispheric uplift -2.3 km tall or >1450 average sized mountains) must be accommodated in ways other than mountain uplift if the lithosphere of lo is 50 kilometers thick, as was suggested by Keszthelyi and McEwen [1997]. Other studies have proposed a lower limit for lo's lithospheric thickness of 30 km [Nash et al.,

1986; Carr et al., 1998]. That would suggest that a volume equivalent to ~450 mountains

(or a hemispheric uplift of -0.7 km) is accommodated in other ways. Lithospheric thicknesses of tens of kilometers are probably reasonable, but the extremely thick (>100 65 km) lithospheres considered in Anderson et al. [2001] and Ross and Schubert [1985] seem difficult to reconcile with the observations and calculations of this study.

2.3.3: Can Hotspot Swells Influence Ionian Mountain Formation?

It is clear that global subsidence and the consequent thermal expansion are able to drive orogenic tectonism on lo. Hovi^ever, these stresses alone may not be sufficient to explain the distribution of ionian moimtains. Turtle et al. [2001] show that, without a focusing mechanism, global compressive stresses acting on an extensively faulted

Uthosphere are likely to produce parallel mountain ranges rather than isolated massifs; thus, there must be a mechanism that localizes compressive stresses. This section discusses the possibility that the dynamics of lo's asthenosphere may play a role in focusing lithospheric stresses, facilitating the uplift of isolated mountains such as are observed on lo.

The thermal structure of lo's interior has been the topic of much discussion. Segatz et al. [1988] present models for two possible end members, one in which lo is largely solid with a thin (50-200 km), partially molten asthenosphere, and the other in which the bulk of lo is partially molten. Some theoretical and observational studies favor a thick (-400 km) asthenosphere with up to 40% partial melt [Spohn, 1997; Keszthelyi et al., 1999,

2004a]. Other observations suggest that the heating and therefore partial melting within lo is concentrated in the upper mantle [Ross et al., 1990; Radebaugh et al., 2001].

Ojakangas and Stevenson [1986] and Greenberg [1982] propose that a -10® yr periodicity in lo's tidal heating causes it to fluctuate between a largely solid and a more molten state. 66 with lo currently in the more molten state. However, Moore [2001] argues that lo cannot transition from a largely solid state to a largely molten one because of negative feedbacks in the tidal heating process. Despite these debates, all models for lo's interior suggest an asthenospheric layer at least 50-200 km thick that is strongly heated by tidal dissipation.

Tackley et al. [2001] model the resulting convection and suggest that the interior of lo is likely to be vigorously convecting. This convection can break up the hemispheric upwelling regions of the Ross et al. [1990] model into numerous smaller plumes. The large number of volcanic centers scattered about lo's surface provides some support for the idea of many local hotspots.

Wherever thermally buoyant asthenospheric diapirs impinge on the base of the hthosphere, the overlying material is thinned and upwarped [e.g., Crough, 1979, 1983;

Sleep, 1992]. So, it seems likely that local topographic swells will form on lo in response to thermal asthenospheric diapirs, and that these swells will alter the stress state of the hthosphere. Theoretically, a lithospheric swell on lo should focus horizontal compressive stresses into an inverted conic shape originating at the base of the hthosphere directly over the diapir head and intersecting the surface at the perimeter of the swell (Fig. 2.10). This predicted stress pattern is corroborated by axially symmetric, two-dimensional finite element modeling that takes into account the surface curvature of lo [Turtle et al., 2001]. In the most simplistic case, failure along a sector of this cone will result in a thrust fault, while failure of the entire cone is predicted to uplift a quasi- circular mountain. The former is thought to be the more likely scenario because, once a thrust fault is initiated, it is easier to continue faulting along that plane of weakness than 67

ippie

c. f.

Figure 2.10. Schematic diagram illustrating the predicted stress distribution and deformation caused by a buoyant diapir impinging on the base of lo's lithosphere (a-c). This is contrasted with the stresses and deformation expected in the absence of a lithospheric horizontal compressive stress (d-f). Stresses are represented as blue arrows in diagrams (a) and (d). The resulting deformation is shown in figures (b) and (e), respectively, in cross section and in figures (c) and (f) in plan view. The stress state that results from an upward force exerted at the base of the lithosphere (d) is coupled with a horizontal compressive stress in diagram (a). While case (d) leads to tensile fracturing of the surface directly over the impinging diapir (e, f), case (a) leads to thrust faulting where the fault plane originates at the top of the diapir and breaches the surface at the perimeter of the swell (b, c). It is apparent in figure (c) that the trace of the thrust fault is slightly arcuate and the edges of the thrust sheet are regions of simple shear offset. 68 to initiate a new fault (c.f Fig. 2.7). By this mechanism, mountains are predicted to form as isolated blocks rather than as parallel ranges.

This model also provides a natural link between mountain and patera formation. On

Earth, an asthenospheric plume impinging on the base of the lithosphere typically produces a volcanic construct at the surface. On lo, the interactions are more complicated. The ascent of magma will be strongly inhibited by the large horizontal compressive stress that acts over much of the lithosphere. Thus, magma should not readily ascend to the surface until thrust faulting has locally relieved the enormous compressive stress induced by subsidence and thermal expansion. Once that stress has been accommodated by mountain uplift, magma supplied by the thermal diapir may buoyantly rise toward the surface using the overlying fractures and faults as conduits for its ascent. Because the thrust fault is a compressive interface, magma may ascend more readily at the tear faults along the margins of the thrust sheet (Fig. 2.10c). Similarly, if thrust faulting over-shortens material in the upper lithosphere as McKinnon et al. [2001] suggest, extension may occur at the trailing edge of a thrust sheet, and paterae may preferentially form there as volcano-tectonic depressions. However, once thrust faulting effectively dissipates the local compressive stress, magma may use the thrust fault plane as a conduit as well.

The focusing mechanism described above is seemingly an inevitable consequence of vigorous convection within lo's asthenosphere and it is consistent with the observed association between mountains and paterae. However, thrust faulting localized by other mechanisms can also explain the moimtain-patera relationship since the fault geometry 69 would be similar. As Turtle et al. [2001] suggest, spatial variations in resurfacing rate and petrologic or structural heterogeneities in the hthosphere are also likely mechanisms for uplifting isolated mountains. The model presented here may play a role in localizing the uplift of some, but certainly not all, of lo's mountains, hi order to identify the processes that drive orogenic tectonism on lo, it is necessary to examine individual mountains and compare their features to those predicted by models.

2.4: Observed Structures

A detailed analysis of individual mountains and their surroundings may shed light on regional tectonic histories. Clues to a mountain's past that are likely to be observable in

Voyager and Galileo data include mountain morphology, any nearby kinematic indicators, and spatial and crosscutting relationships with volcanic features and with other tectonic landforms. Using these data one can compare the observed features with those predicted by the various models. It is also theoretically possible to test mountain formation models by identifying mountains in various stages of uplift and comparing those with predicted deformation sequences. The chief caveat of this investigation is that volcanic and erosional processes rapidly obscure many structural features on lo.

2.4.1: Mountain Morphology

One readily observable feature that models can predict is moimtain morphology.

Thrust faults generally produce asymmetric mountains where one flank dips more shallowly than the other (c.f fault-bend folds), hi the textbook case, the flanks of the 70 mountain have dips that are normal to the strike of the fault, with the steeper flank dipping in the direction that the hanging wall was translated. If mountains on lo form predominantly by thrust faulting, as most models suggest, then many of the less degraded mountains should have this shape. Schenk and Bulmer [1998] showed that East Euboea

Mons has an asymmetric morphology consistent with thrust faulting, and we find that many other mountains display similar characteristics. For example, North Zal Mons is a tilted block with a relatively planar surface that dips gently to the west and drops off abruptly to the east [Turtle et al., 2001]. The mountain's asymmetry suggests that the hanging wall should have moved east. However, the plateau-like appearance of North

Zal Mons is indicative of considerable degradation, which makes structural interpretation difficult. A clearer example is Gish Bar Mons. Like North Zal Mons, it dips steeply to the east and more gently to the west suggesting that the hanging wall moved east.

Shamshu Mons (Fig. 5b) is yet another example. Its southeastern flank has a shallower dip than the northwestern face suggesting a northwestward translation of the hanging wall. Similarly, the mountain next to Shakuru Patera (23.6°N, 269.0°W) has a gently dipping eastern flank and a steeply dipping western flank, which point to a westward translation of the hanging wall.

The longitudinal ridges that sometimes form on mountains may provide a qualitative assessment of slope. Previous studies have suggested that these ridges are compressive folds that form by the downslope movement of a weak layer sliding along a shallow decollement [Heath, 1985; Moore et al., 2001; Turtle et al., 2001]. If this interpretation is correct, the ridges should be oriented such that their fold axes are normal to the dip 71 direction of the decollement. Assuming that the decollement occurs at the boundary between bedrock and an unconsolidated mantling deposit of uniform thickness (e.g., a layer of volatile-rich plume deposits), the dip direction of the decollement and that of the surface should be similar. Relatively few of lo's mountains have been imaged at resolutions where ridges are detectable, but for the majority of mountains with observed ridges, the longitudinal axes of the ridges and the mountain are parallel (e.g., the Hi'iaka

Montes). This observation suggests that the dip directions of those mountains are perpendicular to their long axes [c.f Schenk et al., 2001], which is consistent with a thrust fault origin for the moimtains because the strike of a thrust fault is typically long compared to its heave [e.g., Schultz et al., 2001].

One notable characteristic is that the amplitudes and wavelengths of ridges on individual mountains appear very uniform even though there are obvious differences in ridge morphology from mountain to mountain. This uniformity suggests that there is little variation in the gradients of these mountain flanks and that the massifs were uplifted and tilted as fairly coherent blocks despite the pervasively fractured nature of the ionian lithosphere. In the case of N. Hi'iaka Mons (Fig. 2.5a), stereo data presented by Schenk et al. [2001] show a gently and relatively uniformly sloped plateau. Such evenly tilted planar surfaces are consistent with thrust faulting but could also be formed by extensional or near-vertical tectonics. It should be reiterated that these inductions are based on a series of (reasonable and self-consistent, but unproven) assumptions and more analyses must be completed before conclusive statements can be made. 72

The shape of a mountain's base can sometimes record its movement history. The diapir focusing mechanism discussed in the previous section predicts a circular concentration of compressive stress at the surface. This is likely to produce a thrust fault with an arcuate trace, as is observed at several ionian mountains. However, this fault profile alone is not diagnostic of the diapir model. Crescent shapes are common in thrust faults because they allow for maximum compression at the center of the thrust sheet and progressively less slip towards the margins. Elliot [1976] coined the "bow and arrow rule," which states that, if the trace of a thrust fault is pictured as a bow, the arrow strung on that bow points in the direction that the hanging wall moved. Monan Mons is a perfect example of this geometry (Fig. 2.11). The trace of the thrust fault is concave westward, indicating that the hanging wall moved east. Similarly, the eastern peak in the

Dorian Montes and the mountain chain centered at ~5°S, 318.7°W have crescent shapes that indicate an eastward translation of the hanging wall, and the arcuate shapes of the ridges on Ionian Mons and Mongibello Mons suggest that their hanging walls moved northeast and east, respectively.

2.4.2: Kinematic Indicators

Thrust sheets are often boimded by paired zones of simple shear offset that have shear planes oriented perpendicular to the thrust fault (Fig. 10c). These two shear zones have opposite senses of offset and are typically manifest as tear faults [e.g., Marshank 1988].

They may also have a component of pure shear, and therefore transtensional or transpressional structures may develop. Evidence of tensile deformation (i.e., graben) 73

Figure 2.11. The image on the left shows Monan Patera (top) and Ah Peku Patera (bottom) at either end of the crescent-shaped mountain, Monan Mons (15.4°N, 104.0°W). Monan Patera is a green-floored depression with red diffuse deposits along its margin. Ah Peku Patera as a variegated floor with red, green, orange, and yellow patches. In the schematic diagram on the right, the solid black line is the "bow" that defines the curvature of the mountain, the dashed line connecting its ends is the "bowstring," and the red arrow "strung on the bow" indicates the movement direction of the hanging wall during thrust faulting. The dashed lines and red arrows in the paterae indicate the senses of offset on and approximate positions of the shear zones that bound the thrust sheet. 74 might be found at the traihng edges of the thrust sheets if, as McKinnon et al. [2001] suggest, over-shortening occurs in the upper lithosphere. Graben formed in this manner should be oriented roughly parallel to the thrust fault. Variations in the dip of the thrust fault plane can give rise to complex fold structures in the hanging wall. These folds are potentially observable in the available data after more extensive analysis, and could provide insight into the geometry of the fault plane. To date, kinematic indicators of these sorts have not been directly identified on lo, probably due to the pervasive and continually replenished mantle of volcanic materials and mass wasting deposits. Instead, they have only been suggested indirectly by the positions of volcanic features such as paterae and eruptive fissures.

As mentioned above, faults or fracture zones that are not subject to a large compressive stress are likely conduits for magma ascent. The trailing edge of a thrust sheet is one such region, and the tear faults that bound a thrust sheet are another.

Therefore, paterae may preferentially form at these locations. Shakuru Patera is an elongate, lava-filled depression that appears to lie at the trailing (east) edge of a thrust sheet that moved westward. Its longitudinal axis is parallel to that of the adjacent mountain to the west. Tensile faults formed during mountain uplift may have facilitated the formation of Shakuru Patera by charmeling magma to the region. Paterae also occur along the shear zones that form at the margins of mountains. Monan Patera is an elongate, green-floored depression with active volcanism occurring around its margin as evidenced by red deposits [c.f, McEwen et al., 1998; Geissler et al., 1999; Lopes-Gautier et al., 1999; Phillips, 2000]. Its longitudinal axis is roughly perpendicular to that of 75

Monan Mons. Monan Patera appears to be situated at the edge of a thrust sheet along what is inferred to be a sinistral tear fault (Fig. 11). A second patera (Ah Peku Patera) occurs at the dextral shear zone on the opposite side of the thrust sheet (and at the opposite end of the mountain). The eastern peak of the Dorian Montes has a remarkably similar geometry with paterae at each shear zone bounding the thrust sheet. The paterae at either end of Pan Mensa may also have formed over shear zones, but the sense of offset is unclear in the lower-resolution Voyager data.

Once active thrust faulting ceases, the fault plane itself may also serve as a conduit for magma ascent, hi this case, volcanism would occur, and might form paterae at the leading edge of the thrust sheet along the fault trace. We mention this possibility for the sake of completeness, but find no unambiguous example of a patera located on the trace of a thrust fault, though the best candidate is Zal Patera. The morphology of North Zal

Mons weakly suggests that the thrust sheet moved eastward toward the present location of Zal Patera, a shallow depression partially covered by recent lava flows. The lava flows emanate firom a scarp that runs along the eastern margin of North Zal Mons.

Further to the south, the scarp blends seamlessly into a fracture that runs half way down the western margin of South Zal Mons and terminates at a small patera. This scarp may be the trace of an inactive thrust fault.

2.4.3: Crosscutting Relationships

Where faults and fractures influence the locations of ionian paterae, tectonism must precede volcanism. Therefore, crosscutting relationships can be used to test various 76 hypotheses. There are several examples of mountains incised by paterae: Tohil Mons,

Euxine Mons, Pan Mensa, the Monan Mons, Gish Bar Mons, the low mountain surrounding Chaac Patera, and many others. We infer from the crosscutting relationships that in all of these cases the mountains existed prior to the paterae. This is best imaged at

Tohil Mons (Fig. 2.12). The northeastern flank of Tohil Mons is incised by Radegast

Patera, a narrow, dark-floored depression that lies between the mountain and Tohil Patera

[Turtle et al., 2004]. The southeast margin of Radegast Patera cuts into Tohil Mons forming a tall, steep cliff.

2.4.4: Summary of Observations

The morphology and geologic setting of lo's mountains strongly support a predominantly thrust fault origin and provide ample evidence for a link between tectonism and volcanism. A number of mountains are textbook examples of crustal blocks uplifted along thrust faults. Less decipherable mountain morphologies can be explained by a combination of more complex thrust fault geometries and mountain degradation (primarily mass wasting processes). The locations and relative ages of volcanism in association with mountains are also consistent with a thrust fault model for mountain formation.

Many observations are also consistent with the predictions of the hotspot swell model. However, features diagnostic of this specific mechanism for focusing lithospheric stresses have not been identified. Ideally, one would observe broad (few hundred kilometer scale), low-relief swells at the trailing edges of thrust sheets. Gaskell et al. 77

Figure 2.12. A high-resolution mosaic across Tohil Mons (lower 3 frames) and Tohil Patera (upper 2 frames). The small, dark-floored Radegast Patera between Tohil Mons and Tohil Patera (middle frame) incises the mountain leaving a tall, steep cliff. Tohil Mons is located at 28.9°S, 160.3°W. 78

[1988] reported seeing similar, though larger, features in the Voyager data, but those basins and swells have not been confirmed by Galileo data analysis [Thomas et al.,

1998], Oberst and Schuster [2004] continue to investigate the shape of lo; however, they can only detect (though do not find) very broad swells (on the order of 1000 km). Thus, they probably do not have the resolution to identify the types of swells that this model requires.

2.5: Conclusions

A total of 101 mountains are well enough imaged in the Voyager and Galileo data sets to discern their morphologies and their spatial relationships with surrounding geologic features. The results of this study are similar to those of other studies [Carr et al., 1998; Schenk et al., 2001]. About 4% of these 101 mountains are volcanoes and

~96% are tectonic massifs. A large number (>40 out of 97) of these tectonic massifs are in contact with paterae. This is about 10% more than is expected from a random distribution and the probability of this occurring is estimated to be <1.08%. These statistics suggest a genetic relationship between ionian mountains and paterae. One possible explanation for this association is that faults associated with mountain formation serve as conduits for magma ascent, and so facilitate the formation of paterae around mountains.

Recent models for the formation of ionian mountains differ on the driving mechanism for mountain uplift. Schenk and Buhner [1998] propose that subsidence drives thrust faulting whereas McKinnon et al. [2001] argue that thermal expansion at the base of the 79 crust is potentially a more important factor. This study attempts a quantitative comparison of the two processes and finds that for most cases subsidence is more important, but thermal expansion is a non-negligible term. It becomes dominant only if the lithosphere is exceedingly thin (<12 km) or if the regional resurfacing rate is sustained at a value significantly lower than the estimated global average for ~10^ years.

The calculations used to compare the relative importance of the subsidence and thermal stresses can also be used to provide an estimate of the total number of mountains the net lithospheric shortening can produce. This number is strongly dependent on the thickness of the lithosphere, thus the total number of observed mountains can be used to constrain lithospheric thickness. This approach yields a lower limit of 12 km and suggests that a lithospheric thickness >100 km is unlikely.

The morphologies of many ionian mountains are consistent with formation by thrust faults, as suggested by Schenk and Bulmer [1998]. However, a focusing mechanism is helpful for producing isolated moimtains rather than parallel mountain ranges [Turtle et al., 2001]. This study hypothesizes that lithospheric swells, caused by thermal diapirs impinging on the base of the lithosphere, can focus compressive stresses, enabling the uplift of isolated moimtains. However, it is noted that this is unlikely to be the primary mechanism for focusing mountain formation. A detailed comparison of predicted and observed structural features finds evidence consistent with, but not diagnostic of, this mechanism. CHAPTER 3: A KINEMATIC ANALYSIS OF NEFERTITI CORONA, VENUS

Coronae are quasi-circular volcano-tectonic landforms that appear to be unique to

Venus. They are defined by annular fault and fracture belts -60-2600 km in diameter

[Stofan et al., 1992] that can exhibit either tensile or compressive deformation. The topographic expression of coronae spans the continuum between domes and depressions, and many have raised rims with annular moats. Volcanism is commonly associated with coronae and may be manifest either as central cone fields within the coronae or as lava flows emanating fi-om their annular fractures. More than 360 coronae have been identified on Venus [Stofan et al., 1992], and they have a nonrandom global distribution.

Coronae preferentially occur (1) at elevations near the mean planetary radius, (2) in a cluster at 240°E longitude and near the equator, and (3) in chains where coronae are linked by bands of parallel tectonic lineaments [Squyres et al., 1993].

Coronae are thought to form when diapirs (1) impinge on the base of the lithosphere, doming the overlying material; (2) spread radially, creating a plateau-shaped uplift at the surface; and (3) cool, allowing for gravitational relaxation of the overburden [e.g.,

Squyres et al., 1992]. This hypothesis is consistent with the geomorphic expressions of coronae and related features [e.g., Janes et al., 1992]. The full spectrum of structures predicted by this model, from radially-fractured domes to broad depressions with raised tectonic annuU, is observed on Venus. Transitional structures with both radial and concentric tectonic fabrics exist as well [c.f Fig. 2c in Stofan et al., 1992]. 81

Nefertiti Corona (Fig. 3.1) is a 420 km x 240 km structure located at 36N, 48E in the

Bell Regio quadrangle (V-9). Topographically, it has a raised annulus and a central depression, and it sits atop a local rise (Fig. 3.1b). Lava flows that appear to originate from Nefertiti's aimulus spread out radially around the corona [Campbell and Campbell,

2002]. Synthetic aperture radar (SAR) images collected by the Magellan spacecraft reveal that Nefertiti is a structurally unique corona with a complex kinematic history.

3.1: Kinematic Indicators at Nefertiti Corona

Kinematic indicators are relatively small-scale features that, when examined collectively, enable larger-scale movement history to be reconstructed. Such structures abound in Nefertiti Corona. The locations of select kinematic indicators discussed in the following text are shown in Figure 3.2.

3.1.1: Location A

The west side of Nefertiti's armulus is a regional topographic high that exhibits a

NNW trending tectonic fabric. A magnified view of the westernmost portion of the annulus (Fig. 3.3) reveals parallel paired sets of dark and bright lines (with the former on the west and the latter on the east in left-look SAR) as well as swaths of sub-parallel bright lineations. These are interpreted as graben and fracture systems, respectively.

Many of the graben and fractures are slightly concave to the ENE, making them roughly concentric with respect to the center of the corona. Further to the north and south, the graben and fractures become more linear and assume a radial, rather than a concentric. 82

200 km ^

Figure 3.1. Left-look SAR (a) and topography data (b) obtained by the Magellan spacecraft over Nefertiti Corona, Venus. Both images are to the same scale. In the SAR data (a) darker shades generally indicate smoother terrain; however, topography can modify the radar return such that slopes facing the radar antenna appear relatively bright. The topography data (b) shows elevations as variations in planetary radius. Mean planetary radius for Venus is ~6052 km. Figure 3.2. Nefertiti Corona in left-look SAR. The boxes labeled A-E show the locations of Figures 3.3-3.8, respectively. 84

Figure 3.3. Location A on Figure 3.2 as seen in Magellan left-look SAR. Tensile faults and fractures. Fractures oriented concentric to the center of Nefertiti (which is to the right of this image) appear to crosscut radial graben. The red arrows show the approximate orientation of maximum tensile strain for this region. 85 orientation. This fracture pattern is consistent with corona formation in an approximately east-west oriented extensional regional stress field [Cyr and Melosh, 1993]. In the southeast region of Figure 3.3, a few graben with faults that strike NNE (an approximately radial orientation) are visible. These are interpreted as relatively older structures because they appear to be crosscut by the concentric fractures. This suggests that an earlier episode of radial normal faulting has been overprinted by more recent concentric normal faulting. Thus, these tectonic fabrics record an evolution from tensile hoop stress to tensile radial stress in the western portion of Nefertiti's armulus. This is consistent with the progressive deformation predicted by the diapir model for corona formation [e.g., Squyres et al., 1992] in which early-stage radial normal faulting gives way to concentric faulting as the diapir head spreads laterally.

3.1.2: Location B

The northeastern part of Nefertiti's annulus is composed of a ~50-km-long, sinistral, brittle-ductile, en-echelon shear zone manifest as a parallel array of sigmoidal fractures

(Fig. 3.4). Where deformation is most pronounced along the shear plane, fractures are spaced -300 m to 1.1 km, with a typical spacing of -675 m. The strike of the shear plane is -120° and the accumulated shear strain (y) is 2.0 (standard deviation (sd)=0.2).

Inherent in the computation of y is the assumption that the rotated fractures initially formed as tension gashes oriented 45° to the shear plane. On the north side of the shear plane, several segmented and diffuse bright bands that are oriented -30° to the shear plane cut across the dominant tectonic fabric. These are likely to be a younger generation 86

Figure 3.4. Location B on Figure 3.2 as seen in Magellan left-look SAR. A ~50-kni- long, sinistral, en-echelon array indicative of brittle-ductile deformation. The red arrows show the relative sense of offset. 87 of tension gashes. An alternative interpretation is that the sigmoidal fractures initially formed as antithetic Riedel shear (R') fractures rather than tension gashes. One piece of evidence supporting this interpretation is that the southern tails of the sigmoidal fractures are oriented -60-80° to the shear plane. If these fractures are progressively deformed R' shears, then ^'is 1.6 {sd=Q.2). Depending on whether the sigmoidal fractures formed as tension gashes or as R' shears, the amount of simple shear offset accommodated by the shear zone is either ~11 km or ~7 km, respectively.

2.1.2: Location C

The southeastern portion of the annulus is comprised of a ~90-km-long dextral shear zone (Fig. 3.5) with its shear plane oriented -118°. This shear zone is manifest in two ways: as faint fractures, perhaps R' shears, deflected in a right-lateral direction and as a small en-echelon fault system that steps up to the right. The deflected fractures have a slightly sigmoidal shape and a characteristic spacing along the shear plane of -2.5 km.

Additionally, the tails of these fractures frend 025°; thus they are oriented -90° to the shear plane. The relative positions of the distal ends of the fractures on either side of the shear zone suggest -15 km of relative displacement. This is similar (to within a factor of two) to the displacement computed for the sinistral shear zone on the opposite side of

Nefertiti's annulus in Location B on Figure 3.2. In the same region as the sigmoidal fractures, an even fainter intersecting tectonic fabric with lineaments oriented 015° (or

-100° to the shear plane) can be seen. Crosscutting relationships indicate that the two sets of fractures formed contemporaneously. Drag folds at the intersections of the two 88

Figure 3.5. Location C on Figure 3.2 as seen in Magellan left-look SAR. A ~90-km- long dextral shear zone characterized by deflected fractures and en-echelon faults that step up to the east. The red arrows show the relative sense of offset. A second, fainter set of fi"actures is oriented vertically in this image. Drag folds on the two sets of fractures (see inset) show sinistral offset on the deflected fractures and dextral offset on the vertical fractures. 89 tectonic fabrics indicate that the sigmoidal fractures (with tails oriented 025°) have small amounts of sinistral offset whereas the intersecting fractures (which are oriented 015°) have minor amounts of dextral offset. The sinisfral nature of the sigmoidal fractures is consistent with R' shears.

The en-echelon fault system occurs in the region of maximum displacement along the shear plane. The step-up-to-the-right geometry within a dextral shear zone indicates that the en-echelon faults are compressive in nature.

3.1.4: Location D

A ~140-km-long (in arc length) and ~15-km-wide arcuate fault system outcrops in between the western margins of the two opposite-sensed shear zones (Fig. 3.6). The fault belt forms a regional high with topography sloping down to either side of it. In left-look

SAR, the faults appear to be bright on the west and dark on the east, and in most places the transition from bright to dark is abrupt. This morphology suggests that material is uplifted in a manner consistent with contractional offset (i.e., thrust faulting). The concave-west nature of the faults suggests that material was primarily translated eastward

(e.g., Elliott, 1976; Marshak et al., 1992; Ferill and Groshong, 1993). The direction of offset is estimated to be 112°.

At its center this thrust belt is comprised of approximately 8 distinct ridges, -500 m to 1 km in width, with a characteristic spacing of ~1-2 km. At first glance this appears to be a relatively narrow zone of deformation, but closer scrutiny reveals that compressive faulting extends at least 50 km to either side of the well-defined thrust belt. In the central Figure 3.6. Location D on Figure 3.2 as seen in Magellan left-look SAR. An arcuate, ~140-km-long thrust belt indicating approximately E-W compression (red arrows). The bright lineaments at the top of this image, which crosscut the thrust belt, are the sigmoidal fractures in Figure 3.4. The NE-SW oriented fractures visible between the red arrows are probably the tails of the sigmoidal fractures. The set of subparallel, E-W lineaments on the left-hand side of the image are probably tensile fi-actures. 91 portion of the thrust belt, the ridges are cut by faint fractures oriented -60°. These fractures appear to be the southernmost extensions of the sigmoidal fi-acture tails in the sinistral shear zone at Location B in Figure 3.2; however, the sigmoidal tails are discontinuous between Location B and Location D. At its northern margin, the thrust belt is also crosscut by the aforementioned sigmoidal fractures from Location B. Here a small amount (375 m at most) of dextral offset can be seen where the sigmoidal fractures have cut the thrust faults. This dextral offset is most consistent with the interpretation of the sigmoidally deformed fractures as R' shears.

A set of sub-parallel, arc-normal, E-W oriented, bright fractures extends ~80 km to the west (concave side) of this arcuate thrust belt. Several of the fractures crosscut one another, suggesting that they formed in two or more episodes. At their western ends, the fractures have a characteristic spacing of ~1 km. The spacing increases as elevation increases to the east such that, where they terminate at the thrust belt, the fracture spacing is ~2 km. This fanning eastward geometry in material that moved to the east suggests that the fractures are tensile. Their morphology is reminiscent of longitudinal crevasses in glaciers, which form parallel to the flow vectors in regions of compressive flow and lateral extension. They are also similar to normal faults on the north side of the

Himilayan arc, which again are interpreted to have formed under parallel compression and lateral spreading [e.g., Jackson, 2002]. If the venusian fractures formed via the eastward franslation of material up a west-dipping slope, then it is likely that they too formed parallel to compressive flow and perpendicular to lateral spreading. 3.1.5: Location E

Further to the east is a pair of intersecting tectonic fabrics overprinted by a narrow thrust belt (Fig. 3.7). The two intersecting sets of parallel lineaments do not appear to offset each other suggesting that they formed contemporaneously. Therefore, these are interpreted to be conjugate shear fractures. The bisectrix of the acute angle between conjugate shear fractures gives the orientation of the principal compressive stress (cti). In this case ai is 112°. The acute angle between the fractures is 45-50°. This is consistent with the 50° acute angle reported for conjugate shear fractures at Ki Corona in southeast

Parga Chasma [Willis and Hansen, 1996]. These small acute angles indicate failure in a brittle material. Similarly, the spacing between parallel fractures in the conjugate shears at Nefertiti Corona (~2 km) and at Ki Corona (-1.25 km [Willis and Hansen, 1996]) are comparable.

3.1.6: Location F

Nefertiti's easternmost margin is comprised of a large (-300 km in arc length x -80 km wide) fold-thrust belt (Fig. 3.8). Like the thrust belt at Location D in Figure 3.2, this is interpreted as a series of compressive uplifts based on the bright radar return of the west facing slopes and the darker appearance of the east-facing slopes in left-look SAR.

However, the transition from bright to dark material at the ridge crests is somewhat gradual in places and more abrupt in others, which suggests folding rather than faulting at the surface in some locations. The ridges here are much broader (-5 km) than the ridges in Location D (-1 km), and are highly asymmetric with west flanks being -1 km wide 93

Figure 3.7. Location E on Figure 3.2 as seen in Magellan left-look SAR (left) and a schematic diagram showing tectonic features (right). Conjugate shear fractures (black lines) indicate approximately east-west compression as does the narrow thrust belt (purple lines) that overprints them. The red arrows show the orientation of the compressive stress that formed the shear fractures. 94

Figure 3.8. Location F on Figure 3.2 as seen in Magellan left-look SAR. A ~300-kni- long X ~80-km-wide fold-thrust belt indicating approximately east-west compression. The red arrows show the orientation of this compression. 95 and east flanks ~4 km wide; ridge crest spacing is ~7 km. This asymmetry is most consistent with west vergence on individual thrust faults, which is difficult to reconcile with the larger-scale morphology of the fold-thrust belt. Its concave-west form suggests eastward translation of material, and the direction of maximum compression is estimated to be ~112°.

An unusual characteristic of this fold-thrust belt is that, despite the compressive nature of the ridges, it is topographically low and has a deep, arcuate trough at its center

(Fig. 3.1b). The maximiun elevation contrast at Nefertiti Corona exists between this trough (-6051 km in planetary radius) and the western portion of the annulus shown in

Location A (~6055 km in planetary radius). If the fold-thrust belt formed where the eastward-moving material on its western side transitioned from being the over-riding to the under-riding sheet, it could help explain both the trough and the west vergence on individual faults. Further research is needed to test this hypothesis.

3.2: Kinematic History of Nefertiti Corona

Collectively, the aforementioned kinematic indicators record a deformation history at

Nefertiti Corona that is unlike anything observed elsewhere on Venus. The structures on the western side of Nefertiti are not unusual. They appear to record a transition from early radial normal faulting to later concentric normal faulting, a deformation sequence that is thought to be typical for coronae [e.g., Squyres et al., 1992]. Additionally, the boudinage-like fault and fracture pattern in the west suggests corona formation in an extensional regional stress field [Cyr and Melosh, 1993], which is also quite common. 96

However, the deformation throughout the rest of Nefertiti Corona is unique for Venus.

Numerous kinematic indicators record a self-consistent story in which material in the interior of Nefertiti Corona moved to the east along a vector oriented -115° (Fig. 3.9).

Crosscutting relationships between the fractures in Locations B and D (Figs. 3.2 and 3.6) suggest that the structures are not coeval, but rather deformation propagated from west to east. However, this material acted as a semi-coherent thrust sheet with most of the deformation occurring at its distal end and relatively little interior deformation, although the minor amount of interior deformation that did occur was compressive in nature. The shear zones that border the thrust sheet to the north and south probably accommodated 7-

15 km of simple shear offset. Additional slip may have been accommodated in structures that are no longer visible due to volcanic and/or tectonic resurfacing. This kind of large- scale lateral translation of material in venusian coronae has not been previously reported.

The sigmoidal fractures in the en-echelon array to the north are evidence of localized semi-brittle deformation at the surface of Venus, again something that has not been previously reported for Venus.

3.3: Implications

The extent to which a visco-elastic material such as rock can undergo ductile deformation is a fiinction of both viscosity and strain rate. However, in high viscosity materials such as the dry igneous rocks thought to comprise the crust of Venus, semi- brittle structures like the observed en-echelon array could only form on timescales significantly longer than the age of the venusian surface (-500 Ma [e.g., Strom et al.. 97

Figure 3.9. The kinematic history of Nefertiti Corona as indicated by Figures 3.3-3.8. The red arrows show the inferred movement history. 98

1994]). Mackwell et al. [1998] suggest that dry diabase closely approximates the rheology of the venusian crust, and they give the following flow laws for dry diabase samples from South Carolina and Maryland, respectively:

s = [3.1]

£ = [3.2]

where s is strain rate, cris stress in MPa, R is the universal gas constant, T is temperature in K, the activation energy is in units of kJ/mol, and the preexponential term is in units of

MPa'^^^s"'. Assuming that Venus' crust also has a shear strength similar to that of diabase

(<60 MPa [Attewell and Farmer, 1976]) and that the en-echelon shear zone at Location B

1 c yn (;^1.6-2.0) formed under the maximum possible shear stress, it would take 10 -10 years for this structure to form at the current surface temperature of 750 K. Thus, the semi-brittle deformation at Nefertiti Corona cannot be explained merely by a low strain rate. Instead, the viscosity of the near-surface rock must have been locally reduced. This reduction in viscosity could be either thermally or compositionally induced. As coronae are generally assumed to form over thermal diapirs and are associated with extensive volcanism, a reduction in viscosity due to local heating seems likely; however, a compositional effect cannot be excluded. Assuming the former and using Equations [3.1] and [3.2], the near-surface crustal temperature must have been at least 200 K higher than its present value in order to form the en-echelon shear zone in 500 Ma. 99

Geophysical models that aim to reproduce tectonic deformation at coronae indicate that the thickness of the elastic Uthosphere at the time of radial and annular fracturing was less than -10 km [Janes et al., 1992; Cyr and Melosh, 1993; respectively]. Analyses of gravity field data collected by the Magellan spacecraft suggest current elastic

Uthosphere thicknesses of 15-45 km at coronae and elsewhere on Venus [Sandwell and

Schubert, 1992; Smrekar, 1994; Phillips, 1994]. This difference indicates that either corona formation occurred at an earlier time in Venus' geologic history when the elastic

Uthosphere was globally thinner or the thermal diapirs responsible for corona formation are effective at regionally thinning the elastic Uthosphere. Either possibility could help to explain the en-echelon array at Nefertiti Corona; however, it is also necessary to invoke either additional heating or lower strain rates at Nefertiti since such semi-brittle deformation features are not seen at other coronae. Nevertheless, it seems likely that the en-echelon array developed during corona formation when the elastic Uthosphere was thinner and a diapir supplied the heat necessary to reduce the crustal viscosity. Because the en-echelon shear zone formed in response to the center of Nefertiti Corona moving to the east, it follows that this eastward motion is also likely to have occurred during corona formation. The large-scale eastward translation of material, which propagated from east to west, could have been triggered by (a) asymmetric spreading of the diapir head, (b) structural heterogeneities in the Uthosphere, or (c) gravitational sliding during the doming stage of corona formation. Further work is necessary to discern between these three possibilities. 100

3.4: Conclusions

Kinematic indicators at Nefertiti Corona record an interesting deformation history in which the corona's center slid to the east as a largely coherent thrust sheet with the vast majority of deformation occurring at the eastern margin. The top and bottom of this thrust sheet are bound by sinistral and dextral shear zones that record -7-11 km and -15 km of offset, respectively. The sinistral shear zone in the north is manifest as an en- echelon array formed by semi-brittle deformation, indicating that the viscosity of the near-surface crust was locally reduced at the time this shear zone formed. The reduction of viscosity was probably caused by >200 K of heating due to the presence of an underlying thermal diapir during corona formation. 101

CHAPTER 4: BASALTIC RING STRUCTURES

A condensed version of this chapter is in preparation for submission to the journal

Geology. In order of appearance, the authors on that paper are W. L. Jaeger, V. R. Baker,

A. S. McEwen, L. P. Keszthelyi, D. M. Burr, J. P. Emery and H. Miyamoto. All authors participated in fieldwork in the Channeled Scabland. Devon Burr and Alfred McEwen provided funding for that fieldwork. Devon Burr also provided geologic context for the martian ring structures, analyzed their morphology and aided in their interpretation. Vic

Baker provided a wealth of knowledge about the Pleistocene flooding in eastern

Washington. Laszlo Keszthelyi participated in all three fieldwork sessions and served as the expert on the Columbia River Basalts. Josh Emery aided in the interpretation of the

Banks Lake structure. Hideaki Miyamoto snapped all of the low-altitude aerial photographs. I did all other work, namely collecting and interpreting structural data, reviewing previous models for the origin of the Odessa Craters, conceptualizing the

"phreatovolcanism + inflation" model, mapping lithofacies at the field site near Tokio

Station and interpreting these data, and comparing the terrestrial and martian ring structures.

4.1: Introduction

Basaltic ring structures (BRSs) are enigmatic landforms that were studied in the

1970's and have recently attracted renewed interest after Mars Orbiter Camera (MOC) images revealed morphologically similar structures in Athabasca Valles and Marte 102

Valles, Mars. The BRSs occur in the Channeled Scabland of eastern Washington State, a region where Pleistocene floods from glacial Lake Missoula scoured the surface of the

Columbia Plateau, eroding into the Miocene Columbia River Basalt Group [e.g., Baker

1978]. Because of their unique geologic setting, the study of BRSs can provide insight into two of the most extreme geologic processes that the Earth endures: flood basalt volcanism and catastrophic aqueous flooding.

There are two major populations of BRSs in the Channeled Scabland, the "Odessa

Craters" near Odessa, WA and another unnamed group northeast of Tokio Station, WA

(Fig. 4.1). Although the smaller Tokio BRSs bear more resemblance to the martian structures, the Odessa Craters are better exposed, and therefore have been the focus of more research. This study investigates the formation of the Odessa Craters and the Tokio

BRSs and examines the extent to which either or both are good terrestrial analogs for the martian ring structures (MRSs).

4.2: Field Observations of the Odessa Craters

There are >100 BRSs in the vicinity of Odessa, WA [Hodges, 1978], most of which are clustered along flood-carved valleys [Dayharsh, 1969; McKee and Stradling, 1970].

These "Odessa Craters" occur in the Roza Member of the Wanapum Basalt, which was emplaced ~15 Ma [e.g., McKee and Stradling, 1970]. The structures are defined by quasi-circular annuli 50-500 m in diameter [Hodges, 1978]. Within each annulus, columnar basalt outcrops in a series of concentric, and frequently discontinuous, arcuate ridges and scarps that typically exhibit 1-10 m of vertical relief. Interior topography is 103

Banks Grand Coulee Lake

nJanks Lake Cross Section Wilbur

Davenport Coulee City

• Harrington

X Odessa C I'aters

Odessa Sprague

lokio URSs X Sprague Lake Moses .Lake

Ritzville Moses Lake 20 miles

Figure 4.1. Map showing the locations of field sites in the Channeled Scabland of eastern Washington State. 104 variable; some structures have small central peaks surrounded by moats, but most have central depressions.

This study examined in detail five Odessa Craters: Amphitheater Crater, Cinnamon

Roll Crater, Colosseum Crater, Rock Rose Crater, and Tule Crater (Fig. 4.2, Table 4.1).

(Names in italics are not official.) Each of these has a multi-ringed annulus; two of the structures have central mounds, two have central depressions, and one is relatively planar.

The outcrops at these five Odessa Craters can be divided into three lithofacies: country rock (i.e., the original solid crust of the lava flow), autointrusive dikes (i.e., fed from within the lava flow) and central mound material (Fig. 4.3). The outermost ring typically is comprised of country rock. Further in, the concentric ridges and scarps consist of both country rock and dikes. Frequently, single outcrops will contain a combination of these two lithofacies with the contact exposed, and occasionally, dikes emanating fi-om the interior of the lava flow are preserved in cross section (e.g. the western edge of Tule Crater). Well-formed, narrow (15-30 cm) columns in much of the country rock suggest that it cooled rapidly, perhaps subaqueously [Long and Wood,

1986]. Narrow columns in the dikes also suggest rapid cooling, but this may, in part, be due to their large surface area-to-volume ratio as dikes are typically <2 m wide. The central mounds at Rock Rose Crater and Tule Crater are comprised of irregularly jointed basalt with 20-150-cm-wide, relatively rectilinear columns suggesting subaerial cooling.

The center of Cinnamon Roll Crater (a roughly planar structure), however, is composed of hackly jointed basalt with a high glass content, suggesting very rapid quenching. 105

Figure 4.2. Photographs of the five Odessa Craters examined in this study. Figures a, c and d are oblique, low-altitude airphotos, figure b is a photo taken from the ground and figure e is a USGS digital orthoquad (DOQ). Amphitheater Crater (a) is ~200 m in diameter; south is to the top. Cinnamon Roll Crater (b) is -90 m in diameter and lies on a 14° slope; the photograph looks north. Colosseum Crater (c) is -190 m in diameter; south is to the top. Rock Rose Crater (d) is -160 m in diameter; northwest is to the top. Tule Crater (e) is -270 m in diameter; north is to the top. 106

TABLE 4.1. Locations and sizes of the five Odessa Craters examined in this study.

Odessa Crater Latitude Longitude Diameter (m) Central Topography Amphitheater 47°25.34TSr 118°40.96'W 200 crater Cinnamon Roll 47°24.80'N 118°40.73'W 90 planar Colosseum 47°25.36'N 118°40.52'W 190 crater Rock Rose 47°25.54'N 118°41.41'W 160 peak Tule 47°23.16'N 118°45.25'W 270 peak 107

Figure 4.3. Field sketch of lithofacies at Tule Crater overlaid on USGS DOQ 47118D73. The lithofacies are mapped as follows: purple = country rock, green = autointrusive dikes, and orange = central mound material. While both country rock and dikes exhibit relatively narrow columns indicative of water-cooling, the broad columns of the central mound material are consistent with subaerial cooling. 108

Hodges [1978] reported palagonite in the central mounds of three structures, but palagonite was not seen in the five BRSs examined in this study.

At each of the five structures, the geometric orientations of basalt columns were recorded at numerous locations (Fig. 4.4-4.8; Tables 4.2-4.6). These data show that the

BRSs are approximately axially symmetric landforms with country rock columns plunging steeply (mean=76°, CT=11°) away from the geometric center and dike columns plunging moderately (mean=37°, a=13°) towards the center (Fig. 4.9). Columns in central mounds are approximately vertical. If the dikes cooled from the sides such that their columns are perpendicular to their walls, then, on average, the dikes dip 53° away from the centers of the BRSs. Collectively, these findings indicate that the BRSs are an intrinsic part of the lava flow and that the Missoula floods only served to expose and highhght these preexisting structures.

Four of the BRSs examined in this study are relatively horizontal. In contrast,

Cinnamon Roll Crater lies on a 14° slope. Structural data collected at Cinnamon Roll

Crater was analyzed to determine whether this slope is an inherent part of the structure or an erosional overprint. Two stereonet plots were produced: one showing the trends and plunges of basalt columns as measured from horizontal and the other showing adjusted values taken relative to the 14° slope (Fig. 4.10). The plots show that, when column orientations are considered with respect to horizontal, a more axially symmetric structure emerges. Thus, it is likely that the slope on which Cinnamon Roll Crater is exposed postdates the formation of the BRS, and it suggests that BRSs, like the lava flows that encase them, form in a near-horizontal orientation. 109

Figure 4.4. Locations where data was collected in Amphitheater Crater. This low- altitude airphoto has been rectified such that north is up. The numbers correspond to the field locations listed in Table 4.2. The positions of the field site are approximate as they were transferred fi-om a field sketch to the airphoto. TABLE 4.2. Trends and plunges of basalt columns at Amphitheater Crater.

Location on Figure 4.4 Trend(°) Plunge(°) Lithofacies 1 167 76 Country Rock 2 040 79 Country Rock 3 110 75 Country Rock 4 086 72 Country Rock 5 083 84 Country Rock 6 046 80 Country Rock 7 065 75 Country Rock 8 023 78 Country Rock 9 200 37 Dike 10 008 68 Country Rock 11 193 36 Dike 12 Oil 69 Country Rock 13 340 81 Country Rock 14 146 28 Dike 15 092 32 Dike 16 301 52 Country Rock 17 115 34 Dike 18 245 83 Country Rock 19 284 71 Country Rock 20 238 84 Country Rock 21 091 41 Dike 22 160 80 Country Rock 23 325 34 Dike 24 354 36 Dike 25 049 33 Dike 26 228 30 Dike 27 180 62 Country Rock 28 344 42 Dike 29 195 79 Country Rock 30 231 72 Country Rock Ill

Figure 4.5. Locations where data was collected in Cinnamon Roll Crater. Like Figure 4.1b, this photograph looks north. The numbers correspond to the field locations listed in Table 4.3. The positions of the field site are approximate as they were transferred from a field sketch to the photo. TABLE 4.3. Trends and plunges of basalt columns at Cinnamon Roll Crater.

Location on Figure 4.5 Trend(°) Plunge(°) Lithofacies 1 275 83 Country Rock 2 262 77 Country Rock 3 315 72 Country Rock 4 220 79 Country Rock 5 000 90 Country Rock 6 135 72 Country Rock 7 030 68 Country Rock 8 016 86 Country Rock 9 003 78 Country Rock 10 339 79 Country Rock 11 070 76 Country Rock 12 150 75 Country Rock 13 236 84 Country Rock 14 012 75 Country Rock 15 184 83 Country Rock 16 339 73 Country Rock 17 359 74 Country Rock 18 332 70 Country Rock 19 308 71 Country Rock 20 277 69 Country Rock 21 050 86 Country Rock 22 135 89 Country Rock 23 234 85 Country Rock 24 007 73 Country Rock 25 Oil 88 Country Rock 26 000 90 Country Rock 27 031 89 Country Rock 28 Vesicles similar to those at top of 27 29 081 87 Country Rock 30 351 47 Dike 31 035 65 Dike 32 323 62 Dike 33 317 37 Dike 34 177 66 Country Rock 35 153 58 ? 36 Hackly fracture pattern 37 113 38 Dike 38 068 45 Dike 39 266 85 Country Rock 40 035 45 Dike 41 035 61 Dike 42 170 27 Country Rock 43 188 34 Country Rock ? = Lithofacies could not be identified with certainty. 113

Figure 4.6. Locations where data was collected in Colosseum Crater. In this low- altitude airphoto, south is to the top. The numbers correspond to the field locations listed in Table 4.4. TABLE 4.4. Trends and plunges of basalt columns at Colosseum Crater.

Location on Figure 4.6 Trend(°) Plunge(°) Lithofacies 1 352 55 Dike 2 175 60 Country Rock 3 319 59 Dike 4 318 70 Dike 5 277 31 Dike 6 255 58 Dike la 250 59 Dike 7b 000 84 ? 7c 299 1 Dike 8 025 66 ? 9 099 66 ? 10 060 87 ? 11 093 73 ? 12 337 47 Dike 13 344 45 Dike 14 347 48 Dike 15 305 43 Dike 16 298 56 Dike 17 162 82 Country Rock 18 070 71 Country Rock 19 Magnetized rock; can't take measurement 20 072 83 Country Rock 21 041 82 Country Rock 22 118 86 Country Rock 23 016 75 Country Rock 24 040 72 Country Rock 25 215 8 Dike 26 200 25 Dike 27 176 20 Dike 28 039 63 Country Rock 29 039 75 Country Rock 30 208 21 Dike 31 213 35 Dike 32 255 43 Dike 33 234 46 Dike 34 217 77 Country Rock 35 240 75 Country Rock 36 000 90 Country Rock 37 198 75 Country Rock 38 044 85 Country Rock 39 000 90 Country Rock 40 225 68 Country Rock 41 039 30 Dike 42 012 33 Dike 43 205 65 Country Rock 44 357 66 Dike 115

TABLE 4.4 - continued

Location on Figure 4.6 Trend(°) Plunge(°) Lithofacies 45 188 83 Country Rock 46 025 44 Dike 47 024 30 Dike 48 313 81 Country Rock 49 333 47 Dike 50 329 50 Dike 51 315 38 Dike 52 300 43 Dike 53 270 51 Dike 54 037 30 Dike 55 066 52 Dike 56 230 78 Country Rock 57 235 79 Country Rock 58 241 80 Country Rock 59 262 79 Country Rock 60 050 32 Dike 61 050 38 Dike 62 063 30 Dike 63 055 20 Dike 64 077 31 Dike 65 074 39 Dike 66 098 16 Dike 67 259 80 Country Rock 68 115 14 Dike 69 284 80 Country Rock 70 289 83 Country Rock 71 284 82 Country Rock 72 200 87 Country Rock 73 113 29 Dike 74 133 30 Dike 75 315 77 Country Rock 76 145 58 Dike 77 145 35 Dike 78 133 45 Dike 79 140 22 Dike 80 128 27 Dike 81 055 30 Dike 82 058 39 Dike 83 050 35 Dike 84 050 66 ? 85 026 54 ? ? = Lithofacies could not be identified with certainty. 116

Figure 4.7. Locations where data was collected in Rock Rose Crater. This low-altitude airphoto has been rectified such that north is up. The poor image quality is due to the oblique angle and low resolution of the original photograph. The numbers correspond to the field locations listed in Table 4.5. TABLE 4.5. Trends and plunges of basalt columns at Rock Rose Crater.

Location on Figure 4.7 Trend(°) Plunge(°) Lithofacies 1 300 81 Country Rock 2 000 90 Country Rock 3 315 84 Country Rock 4 264 24 Dike 5 251 83 Country Rock 6 218 48 Dike 7 233 36 Dike 8 286 57 Dike 9 069 69 Central Moiuid 10 103 69 Central Mound 11 025 87 Central Mound 12 211 86 Central Moimd 13 340 89 Central Moimd 14 351 45 Dike 15 004 36 Dike 16 351 43 Dike 17 140 76 Country Rock 18 005 59 Dike 19 030 76 ? 20 210 82 Country Rock 21 112 40 Dike 22 009 70 Country Rock 23 288 88 Country Rock 24 335 84 Country Rock 25 340 82 Country Rock 26 004 88 Country Rock 27 006 77 Country Rock 28 050 79 Country Rock ? = Lithofacies could not be identified with certainty. 118

Figure 4.8. Field sketch plotting locations where data was collected in Tule Crater. The numbers correspond to the field locations listed in Table 4.6. The sketch is roughly oriented with north to the top. 119

TABLE 4.6. Trends and plunges of basalt columns at Tule Crater.

Location on Figure 4.8 Trend(°) Plunge(°) Lithofacies 1 261 31 Dike 2 283 32 Dike 3 270 49 Dike 4 305 29 Dike 5 299 36 Dike 6 325 31 Dike 7 271 32 Dike 8 280 44 Dike 9a 246 79 a 9b 232 75 a 9c 303 38 Dike 10 287 22 Dike 11 321 25 Dike 12 287 84 Country Rock 13a 090 83 Country Rock 13b 095 59 ? 14 086 86 Country Rock 15 024 85 Country Rock 16 145 81 Country Rock 17 166 74 Country Rock 18 151 73 Country Rock 19 148 85 Country Rock 20 155 77 Coimtry Rock 21 160 72 Country Rock 22 100 50 ? 23 322 28 Dike 24 304 38 Dike 25 126 75 Country Rock 26a 242 89 Country Rock 26b 165 78 Country Rock 27 338 20 Dike 28 350 33 Dike 29 163 85 Country Rock 30 168 80 Country Rock 31 331 70 Country Rock 32 285 74 Country Rock 33 163 72 Country Rock 34 316 44 Dike 35 028 19 Dike 36 181 69 Country Rock 37 178 63 Country Rock 38 140 71 Country Rock 39 162 61 Country Rock 40 171 53 Country Rock 41 320 21 Dike 42 339 22 Dike TABLE 4.6 - continued

Location on Figure 4.8 Trend(°) Plunge(°) Lithofacies 43 163 65 Country Rock 44 347 53 ? 45 351 16 Dike 46 310 44 Dike 47 008 22 Dike 48 028 30 Dike 49 220 54 Country Rock 50 225 70 Country Rock 51 204 50 Country Rock 52 201 52 Country Rock 53 183 57 Country Rock 54 228 70 Country Rock 55a 206 62 P 55b 148 60 P 55c 345 52 P 56 229 57 ? 57 242 39 ? 58a 263 43 ? 58b 255 59 ? 59 261 62 Country Rock 60a 147 33 y 60b 162 70 y 61 236 79 Country Rock 62 290 75 Country Rock 63 249 67 y 64 090 43 y 65 285 54 y 66 255 65 Country Rock 67 254 68 Country Rock 68 164 54 ? 69 157 25 Dike 70 281 55 Country Rock 71 000 90 Country Rock 72 100 38 Dike 73 116 31 Dike 74 092 24 Dike 75 151 34 Dike 76 313 43 ? 77 140 45 ? 78 314 50 ? 79 323 65 ? 80 154 31 Dike 81 324 49 Dike 82 002 62 ? 83 185 25 Dike 84 354 47 Dike 85 181 48 Dike 86 214 42 Dike 121

TABLE 4.6 - continued

Location on Figure 4.8 Trend(°) Plunge(°) Lithofacies 87 219 39 Dike 88 221 26 Dike 89 Subaerially cooled basalt; no jointing 90 173 90 Central Mound 91 271 88 Central Mound 92 039 83 Central Mound 93 053 58 Country Rock 94 309 77 Country Rock 95 124 83 Country Rock a = Fanning joints. P = Dike encasing a sliver of country rock. Y = Fanning columns in autointruding dikes. ? = Lithofacies could not be identified with certainty. 122

Figure 4.9. Map of structural data collected at Colosseum Crater. The white arrows show the trends of basalt columns within the country rock (mean plunge = 77°) and the black arrows show the trends of basalt columns within the dikes (mean plunge = 39°). 123

Figure 4.10. Stereonet plots showing trends and plunges of basalt columns at Cinnamon Roll Crater. Figure a shows the original data measured with respect to horizontal. Figure b shows the same data measured with respect to the 14° slope on which Cinnamon Roll Crater is situated. The data in plot a clusters more evenly around the center of the stereonet (+). This suggests that Cinnamon Roll Crater formed in a horizontal orientation. 124

4.3: Banks Lake Cross Section

McKee and Stradling [1970] described what they interpreted to be an Odessa Crater­ like structure exposed in cross section on the west wall of upper Grand Coulee over

Banks Lake (Fig. 4.1). However, Self et al. [1996] interpret the same structure as the contact between two inflated pahoehoe sheet flows. Because of its relatively inaccessible location, both McKee and Stradling [1970] and Self et al. [1996] examined this feature from the opposite shore of Banks Lake at a distance >2 km.

In an effort to better understand the origin of the Odessa Craters, the Banks Lake structure (Fig. 4.11) was examined in situ. This structure is characterized by a broad trumpet-shaped depression (-500 m wide, ~55 m deep) within a 65-m-thick N2 Grande

Ronde Basalt. The depression occurs in entablature with columns ~30 cm wide. Near the center of the depression, the columns plunge at moderate angles radially outward; the plunges progressively steepen, and eventually become vertical once out of the depression.

Within the entablature, 1-2-m-wide dikes outcrop at radial distances of 15-100 m. Three dikes outcrop in the northern half of the structure and at least one outcrops on the south side. The dikes are all concave with respect to the center of the depression, and where they intersect the walls of the depression, they dip ~30° away from the geometric center of the structure. The dikes do not have clear terminations at either end. Their bases emanate from the lower portion of the lava flow, indicating that they are autointruded, and their tops appear to have fed the lava that fills the depression.

At the center of the depression and near the base of the lava flow, there is a small

(~2.5 m X 6 m) outcrop of highly altered, 3-10 cm, subrounded, subequant lava clasts 125

Figure 4.11. An Odessa Crater-like structure exposed in cross section on the west wall of Banks Lake, WA, at 47°48'32.8"N, 119°12'25.8"W. The relevant lava flow, a 65-ni- thick N2 Grande Ronde Basalt, is outlined in black. Near the base of the flow, a mound of partially palagonatized, brecciated basalt with a small amount of spatter is highlighted in yellow. Autointmsive dikes are shown in green. Although it is difficult to discern fi-om this distance, the ends of the dikes blend into the surrounding rock rather than terminating abruptly. The boundary between the original flow surface and the dike-fed lavas that fill the depression is outlined in purple. This "contact" is internal to the lava flow. The two insets at the bottom of this figure show magnified views of the outcrop. Both photographs were taken looking up, so they are foreshortened. The inset in the lower left shows the brecciated mound, which is interpreted to be a rootless cone, and the inset in the lower right shows the northernmost dike. 126 residing in a sparse matrix of sand-sized, palagonatized glass fragments. Many clasts have stretched pahoehoe-like skins and a few retain ropy textures. A small number of clasts have morphologies suggestive of ribbon spatter or bombs. This portion of the outcrop is interpreted to be a small rootless cone (i.e., pseudocrater), which formed as the result of weak steam explosions.

The dike-fed lavas that fill the depression are largely columnarly jointed stacks of horizontal lava flows. In the lowermost 4 m of the depression, however, the lava morphology is more heterogeneous. Just above the contact with the palagonatized breccia, the dike-fed lava contains numerous irregular, vertically elongated pods of vesicular lava. These suggest that minor amoimts of steam entered the lava immediately overlying the central rootless cone.

The Banks Lake structure is similar to the Odessa Craters in many ways. It is comparable in size, lithofacies, autointrusive dike positions, and orientations of basalt columns. Unfortunately, a more quantitative comparison is not possible because the cliff that cuts through the Banks Lake structure may not perfectly bisect it. Thus, the measured dimensions are lower limits and measiu"ed dips may not be true dips. However, the numerous qualitative similarities indicate that McKee and Stradling's [1970] characterization of the Banks Lake structure as an Odessa Crater-like BRS was correct.

The only difference is that, while the Missoula Floods unroofed the Odessa Craters exposing them in plan view, the same floods gave cross-sectional exposure to the Banks

Lake structure. This good fortune provides researchers with a three-dimensional view of these remarkable features. 127

4.4; The Origin of the Odessa Craters

Since the first published description of the Odessa Craters [Grolier, 1965], researchers have proposed several competing hypotheses for their origin. Most of these models involve point-source collapse. Dayharsh [1969] suggested that the collapse was due to removal of support at depth, perhaps via evacuation of a magma chamber.

Alternatively, Parks and Banami [1971] interpreted small, negative (-1 mgal) gravity anomalies as evidence for volcanic pipes under the Odessa Craters and suggested that the volimie loss associated with the solidification of magma in the pipes could produce the collapse structures observed at the surface. However, the autointrusive nature of the dikes invalidates these hypotheses.

After a detailed investigation of the Odessa Craters, McKee and Stradling [1970] proposed the "sag flowout" model (Fig. 4.12a) for their origin, which involves the following sequence of events: (1) a point-source crack develops in the solid crust of a cooling lava flow, (2) molten lava exploits the crack and ponds on the surface, (3) the solid crust sags in response to the loading of the surface and the withdrawal of support,

(4) as the crust sags, concentric fractures form allowing fluid lava from the interior of the flow to escape to the surface, and (5) steps 3 and 4 are repeated (a variable number of times) such that the structure progressively expands outward as it forms. This model elegantly produces the general shape of the Odessa Craters; however, it is inconsistent with the rootless cone in the center of the Banks Lake structure and with the palagonite that Hodges [1978] observed. Figure 4.12. Schematic diagrams illustrating the following formation models for the Odessa Craters: (a) the "sagflowout" model (after McKee and Stradling, 1970), (b) the eroded pseudocrater model (after Hodges, 1978), and (c) the "phreatovolcanism + inflation" model proposed in this study. In all diagrams, medium gray indicates fluid lava and light gray (best seen in frame 3 of figure b) represents water.

K) 00 129

A very different model for the origin of the Odessa Craters was proposed by Hodges

[1978]. According to this model (Fig. 4.12b), the BRSs formed as eroded pseudocraters in the following way: (1) lava filled a preexisting basin in the Odessa area; (2) as the lava cooled and solidified, the water table rose such that it intersected the molten interior of the lava flow where it filled the basin; (3) the interaction of water and lava caused explosive steam venting as well as doming and cracking of the crust; (4) the solid crust subsided around explosive vents and fluid lava intruded into concentric fi-actures; and (5) the Missoula Floods eroded the pseudocraters, shaping the present landscape. Again, this model accurately produces many of the characteristics of the Odessa Craters; however, the position of the small, relatively coherent rootless cone near the bottom of the lava flow is inconsistent with BRS formation at the surface of a stagnated and solidifying lava flow. Additionally, multiple, concentric, autointrusive dikes have never been reported around rootless cones.

The findings of this study are not fully consistent with any of the previous formation models for the Odessa Craters; therefore, we propose a new model for their origin - the

"phreatovolcanism + inflation" model (Fig. 4.12c). By this model, a BRS forms when

(1) a relatively thin lava flow advances over a water-saturated substrate; (2) minor explosive activity disrupts the lava flow and builds a rootless cone; (3) the lava flow thickens by the process of inflation, but inflation is locally impeded by rapid solidification at the rootless cone, and this leads to the trumpet-shaped depression; (4) as inflation continues, tensile radial stresses build around the depression and concentric fi-actures develop; (5) fluid lava from the interior of the flow exploits the fractures and fills the depression; and (6) subsequent erosive floods expose the structure. The steps involved in this model have been abbreviated for the sake of clarity. However, it should be noted that during the continuous process of lava flow inflation, steps 4 and 5 will occur repeatedly. There is also evidence that, during steps 3 and 4, steam generated in the substrate can continue to vent through the path of least resistance (i.e., the rootless cone).

4.5: Basaltic Ring Structures Near Tokio Station, WA

A second population of BRSs can be found northeast of Tokio Station, WA

[Dayharsh, 1969; McKee and Stradling, 1970; Baker, 1978b] (Fig. 4.1). The Tokio structures are distinct from the Odessa Craters in that they are smaller (40-150 m in diameter), more randomly distributed, single-ringed, raised-rimmed structures with little topographic expression and significant loess infill. The specific lava flow in which the

Tokio BRSs occur was not determined, but it is unlikely to be the Roza Member because it lacks the characteristic plagioclase pheoncrysts. Dikes do not outcrop in the Tokio structures; however, it is not known whether they are absent or buried. The single rings that define the Tokio structures are invariably composed of narrow (~15 cm), slightly curvilinear basalt columns that plunge steeply away from the centers of the rings. The

"matrix" that outcrops between the rings is composed of wide (~75 cm), vertical basalt columns. These findings suggest that the Tokio BRSs are sites of localized water-cooling of the lava (Fig. 4.13). Building on that interpretation, the radial outward trends of the basalt columns in the rings may result from (a) rapid cooling catalyzed by small. 131

Figure 4.13. Field map of lithofacies present in the BRSs near Tokio Station, WA. Blue indicates narrow (-15 cm), steeply plunging basalt columns. Orange represents broad (~75 cm), vertical columns. This map suggests that the Tokio BRSs were sites of localized, rapid cooling of the lava. 132 localized water sources at the top of the lava flow, (b) subsidence under rootless cones, as

Hodges [1978] hypothesized, or (c) lava flow inflation around rootless cones. The latter two hypotheses seem more likely because it is difficult to explain how water could cool the top of the lava flow in the observed pattern.

While there are significant differences in diameter and spatial distribution between the Tpkio BRSs and the Odessa Craters, it is still plausible that they formed by a similar process. Pre-lava emplacement topography, groundwater distribution, the extent to which inflation thickens the lava flow, and differential erosion are all factors that could influence the size and/or distribution of the BRSs. Thus, a phreatovolcanic component to the formation of the Tokio BRSs is likely, but this study cannot distinguish between flood-eroded rootless cones or an Odessa Crater-like origin.

4.6: Comparison to Features in Athabasca Valles, Mars

Numerous ring structures that are similar in appearance to the Tokio BRSs are exposed on the floor of Athabasca Valles, a young, low-latitude (~10°N, 155°E) martian channel system that is incised into the Cerberus plains (Fig. 4.14). Studies generally indicate that the geologic history of Athabasca Valles is similar to that of the Channeled

Scabland. The region is covered by flood lavas [e.g., Plescia 1990, 2003; Keszthelyi et al., in press; Lanagan 2004] and has been carved by catastrophic aqueous floods [e.g..

Burr et al., 2002a,b; Berman and Hartmann, 2002; Plescia, 2003]. The lithostratigraphic unit that covers much of the channel floor has been hypothesized to be either an ice-rich mud flow [Rice et al. 2002] or a lava flow [Lanagan et al., 2001; Keszthelyi et al.. 133

Figure 4.14. Comparison of (a) BRSs near Tokio, WA and (b, c) MRSs in Athabasca Valles, Mars. The 200-m scale bar applies to (a) and (b). Image (a) is a low-altitude airphoto reprojected onto USGS DOQ 47118B21 such that north is up, (b) is an excerpt from MOC image ElO-01384 (resolution = 3.09 m/pix), and (c) is an excerpt from MOC image R12-03203 (resolution = 1.53 m/pix) showing three MRSs that appear to be partially buried by (or partially exhumed from) an overlying deposit that Burr et al. (2002) interpret to be a subaqueous dune field. 134

2004a], the latter being more consistent with the observation that the region is rough at radar wavelengths [Lanagan, 2004], The martian ring structures (MRSs) occur in a subsection of this unit that is characterized by patterned terrain. Elsewhere in Athabasca

Valles, the patterned terrain was found to have a relatively high thermal inertia, which is consistent with crystalline rock [Lanagan, 2004]. Because the unit containing the MRSs fills the valley floor, it must postdate the channel-cutting phase of the floods. However,

MOC images ElO-01384, ROl-00745 and R12-03203 show several MRSs that appear to be partially obscured by an overlying deposit that Burr et al. [2002b] interpret to be a field of subaqueous dunes (Fig. 4.14c). This stratigraphic relationship suggests that the

MRSs predate the dunes. Additionally, Lanagan [2004] finds geomorphic evidence that floods eroded young lavas in Athabasca Valles. Thus, it is likely that the MRSs and the terrestrial BRSs occur in similar geologic settings.

Burr et al. [in review] postulate that the MRSs are pingos, collapsing pingos, and pingo scars; however, this interpretation is difficult to reconcile with ice-stability models for Mars, which predict an absence of near-surface ice at equatorial latitudes [e.g.,

Mellon and Jakosky, 1995], and with Gamma Ray Spectrometer data that show this to be one of the driest regions of Mars [e.g., Boynton et al., 2002]. Instead, I suggest that the

MRSs formed in a manner similar to terrestrial BRSs: through a combination of water- lava interactions and flood erosion. This hypothesis is consistent with both the geologic setting and the morphologies of MRSs. A quantitative comparison of BRSs and MRSs

(Table 4.7) reveals that the Tokio and martian populations are fairly similar in diameter, ellipticity and number density. The Odessa Craters are quite different from both the TABLE 4.7. Comparison of BRSs in the Channeled Scabland and MRSs.

Population of rings Mean diameter (m) Ellipticity* Number density (km-') Odessa, WA 200±99^ 1.09±0.1^ 3 Tokio, WA 75±26^ 1.44±0.3+ 26 Athabasca Vallis, Mars 60±29t 1.74±0.7^ 35 *Ellipticity = major axis/perpendicular axis. ^Standard deviation. 136

Tokio and martian rings in that they are larger, more circular and more dispersed.

(Again, this difference could, in theory, arise from variations in factors such as groundwater supply, lava flux, initial topography, and degree of flood erosion.) If the

MRSs are analogous to the Tokio BRSs, then it indicates that groundwater (or, perhaps, ground ice) was present at the time of lava flow emplacement, and it requires at least two floods in Athabasca Valles - one to cut the main charmel and a second to erode the surface of the lava flow and deposit the dunes.

4.7: Conclusions

The "phreatovolcanism + inflation" model is proposed for the formation of the

Odessa Craters. This model postulates that the structures originally formed when Roza lava flows inflated around rootless cones, but their characteristic crater-like morphologies are the result of flood erosion. An important consequence of this model is that the

Odessa Craters are diagnostic of inflated lava flows, water-lava interactions, and subsequent catastrophic floods. Like the Odessa Craters, the Tokio BRSs probably formed around rootless cones; however, it is not clear whether inflation also played a role in their origin. Nevertheless, Tokio-like BRSs are indicative of water-lava interactions and flood-erosion. Similarities between the MRSs and the Tokio BRSs are striking. If future research confirms that they formed by analogous processes, then MRSs may prove to be a useful tool for identifying regions of Mars where both water-lava interactions and flood erosion have taken place. The upcoming Mars Reconnaissance Orbiter (MRO) mission, which is scheduled for launch in August of 2005, may provide the data 137 necessary to more fully understand the formation of the MRSs. The High Resolution

Imaging Science Experiment (HiRISE) onboard MRO should return images with sufficient resolution to allow researchers to distinguish between the pingo and BRS hypotheses for MRSs. 138

CHAPTER 5: SUMMARY

In this thesis I examined a number of planetary bodies using structural geology techniques at a wide range of scales. The primary conclusions from these different studies are:

1) The mountains on lo are primarily uplifted by thrust faulting.

2) The uplift is driven by a combination of subsidence and thermal stresses, but

subsidence stresses should dominate.

3) lo's lithosphere is thicker than 12 km and probably significantly thinner than 100

km.

4) Tectonism on lo plays an important role in localizing where magma can ascend

through the lithosphere.

5) The center of Nefertiti Corona, Venus, was translated to the east.

6) Part of Nefertiti's tectonic annulus experienced semi-brittle deformation, probably

because heating during corona formation reduced the viscosity of the crust.

7) The orientation of cooling joints in the Odessa Craters in eastern Washington

State show that these structures are internal to the Roza lava flow and were

exposed, rather than created, by the Pleistocene Missoula Floods.

8) The presence of spatter and other phreatovolcanic materials internal to, but near

the bottom of, a basaltic ring structure exposed in cross section at Banks Lake is

inconsistent with any of the previously proposed formation mechanisms for the

Odessa Craters. 139

9) Lava inflation around a zone chilled by phreatovolcanic activity provides the best

explanation for the origin of the Odessa Craters.

10) A similar (but not necessarily identical) process is likely to have formed the

basaltic ring structures near Tokio Station, WA.

11) Ring structures on the floor of Athabasca Valles, Mars, are similar in appearance,

diameter, ellipticity and number density to the Tokio rings.

12) If the Tokio and martian rings formed by the same mechanism, it is evidence for

water-lava interactions and at least two floods in Athabasca Valles. 140

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