The Influence of Lithospheric Flexure Induced by Volcano Loading on Neogene Basin Evolution in McMurdo Sound, West Antarctica

THESIS

Presented in Partial Fulfillment of the Requirements for the Degree Master of Science in the Graduate School of The Ohio State University

By

Jie Chen

Graduate Program in Geological Sciences

The Ohio State University

2015

Master's Examination Committee:

Dr. Terry Wilson, Advisor

Dr. Michael Bevis

Dr. Derek Sawyer

Copyright by

Jie Chen

2015

Abstract

The marine Erebus Basin surrounds the volcanic Ross Island in McMurdo Sound,

Antarctica. This basin has evolved under the influence of flexure driven by loading by the three volcanoes that make up Ross Island. Seismic reflection data west of Ross Island are used to document the seismic stratigraphic framework of the flexural basin. Five seismic units are mapped within the sedimentary infill of the basin. Seismic facies are documented within each unit, ranging from chaotic facies interpreted as volcanic mass flows and ice-proximal glacial deposits, to laminated units interpreted as alternating pelagic, hemipelagic and glaciomarine deposits. The volcanogenic units form part of the aprons around the volcano slopes and the glaciomarine units fill the axial basin.

Two seismic sequences are defined based on patterns of surface depth and unit thickness. Seismic Sequence 1 is bounded by the Rj and Rk seismic surfaces, showing thickening and deepening toward Mount Bird volcano. Seismic Sequence 2 includes units from the Rk seismic surface up to the seafloor, and thickens and deepens toward Mount

Erebus volcano. Seismic mapping thus shows evidence of two discrete sub-basins formed by Mt. Bird loading since 3.8-4.6 Ma and Mt. Erebus loading since ~1.31 Ma.

Both units also show thickening in a zone northwest of Ross Island where accommodation space was formed by Terror Rift faulting, indicating that Terror Rift was still active when the Ross Island volcano loading occurred.

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The Bird-Erebus flexural basin has an overall wedge-shaped fill that thickens toward the volcanoes of Ross Island, typical of the geometry of other basins formed by volcano loads in ocean basins. The Erebus Basin is underfilled, with only ~325 m of infill, compared with >2000 m of fill typically found in moats around ocean island volcanic chains. Like stratal sequences of other flexural basins, the Bird-Erebus flexural basin fill shows onlap geometries in cross-moat profiles. However, unlike stratal sequences of other flexural basins, no offlap pattern is observed in either cross-moat or along moat profiles. Gravity flows related to mass wasting of Ross Island volcanoes are a significant component of the Bird-Erebus basin fill but, unlike other basins formed by volcano loads, is not the dominant fill. Instead, seismic facies mapping and AND-1B core show that a significant component of glaciomarine deposits occurs in the flexural basin.

Glacial erosion removed some of the flexural basin fill, particularly along the regional Rk seismic surface.

A 3D thin elastic plate model was used to simulate flexural basin evolution induced by Ross Island volcano loading, and to evaluate regional lithospheric strength.

Using the dip angle of seismically mapped reflector surfaces as the constraint, the best fit lithospheric strength is represented by effective elastic thickness of 4-5 km. This weak lithosphere in the Ross Island region can be explained by faulting, thin crust, and high heat flow in this rift tectonic setting.

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Dedication

This document is dedicated to my family.

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Acknowledgments

Firstly, I would like to express my sincere gratitude to my advisor, Dr. Terry

Wilson, for her support and guidance of my Master study, for her patience, great vision, incredible motivation, as well as the opportunity to do field work in Antarctica. Besides my advisor, I would like to thank Dr. Michael Bevis and Dr. Derek Sawyer for their insightful input from varies perspectives. Also, I would like to thank Dr. Stuart Henrys for his help in seismic work.

The primary marine geophysical data used in this research were acquired on

NBP0401 geophysical cruise, funded by the National Science Foundation. Also, the support of GRA throughout my years at OSU is funded by the National Science

Foundation. An Educational Grant from Schlumberger provided the software Petrel for seismic interpretation.

I would also like to thank members of our research team: Stephanie Konfal, Dave

Saddler, Joel Main, Cristina Millan, Jamey Stutz, Tricia Hall, and Will Blocher. We’ve enjoyed a lot of moments, from playing football at sun shine Antarctica to fighting the deadlines at the basement of Orton. Lastly, I would like to thank my wife and best friend,

Xiaorui. She is always sitting next to me.

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Vita

2003...... Huangyan High School

2007...... B.S. Earth Sciences, Zhejiang University

2012 to present ...... Graduate Research Associate, School of

Earth Sciences, The Ohio State University

Fields of Study

Major Field: Geological Sciences

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Table of Contents

Abstract ...... ii

Dedication ...... iv

Acknowledgments...... v

Vita ...... vi

List of Tables ...... xi

List of Figures ...... xii

Chapter 1: Introduction ...... 1

Chapter 2: Literature Review ...... 3

2.1 Geological Setting ...... 3

2.1.1 The West Antarctic Rift System ...... 3

2.1.2 The Victoria Land Basin and Terror Rift ...... 4

2.1.3 Ross Island Volcanic Complex...... 4

2.1.4 Regional Lithospheric Strength ...... 5

2.2 Isostasy and Lithospheric Flexure Model ...... 7

2.3 Previous Flexural Studies of Ross Island Loading ...... 8

Chapter 3: Data ...... 10

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3.1 Digital Elevation Model (DEM) ...... 10

3.1.1 Land Surface Elevation of Ross Island...... 10

3.1.2 Bathymetric Data for the Seafloor Surrounding Ross Island ...... 11

3.1.3 Integrated DEM ...... 12

3.2 Seismic Reflection Data ...... 13

3.2.1 Reflection Seismology ...... 13

3.2.2 Seismic Data Sources ...... 14

3.2.3 Mis-tie Analysis ...... 15

Chapter 4: Methods ...... 16

4.1 Workflow for this Study ...... 16

4.2 Seismic Interpretation ...... 17

4.2.1 Seismic Stratigraphic Techniques ...... 17

4.2.2 Artifacts in Seismic Data ...... 19

4.2.3 Seismic Interpretation in Petrel ...... 20

4.3 Modeling Method ...... 21

Chapter 5: Results ...... 23

5.1 Seismic Framework ...... 23

5.1.1 Seismic Facies ...... 23

5.1.2 Description of Seismic Surfaces and Units ...... 31

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5.1.3 Age Constraints for Seismic Surfaces and Units ...... 39

5.2 Modelling ...... 44

5.2.1 Velocity Model ...... 44

5.2.2 Model Input ...... 45

5.2.3 Model Output ...... 47

5.2.4 Best Fit Model ...... 48

Chapter 6 Discussion ...... 50

6.1 Flexural Basin Evolution ...... 50

6.1.1 Sequential Flexural Basins Related to Mt Bird and Mt Erebus Volcano Loading

...... 50

6.1.2 Relation to Terror Rift ...... 54

6.1.3 Infilling of Flexural Basin ...... 54

6.1.4 Basin Fill Geometry and Progressive Volcano Loading ...... 58

6.2 Flexural Modeling and Lithosphere Strength ...... 59

6.2.1 Comparison with Previous Ross Island Flexural Modeling ...... 59

6.2.2 Flexural Modeling Results Compared with Other Basins Formed by Volcanic

Island Loading ...... 62

Chapter 7: Conclusions ...... 63

References ...... 66

ix

Appendix A: Tables ...... 73

Appendix B: Figures ...... 78

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List of Tables

Table 1: Seismic stratigraphic framework for this study...... 73

Table 2: List of 7 seismic facies with their reflection characteristics and possible interpretations...... 74

Table 3: Volume of Ross Island volcanos...... 75

Table 4: Parameters used in flexural model...... 76

Table 5: Dip angle of seismic reflector Rj in 20 seismic profiles...... 77

xi

List of Figures

Figure 1. Geological setting of Ross Island ...... 78

Figure 2. Bedrock topography of McMurdo Sound with seismic lines ...... 79

Figure 3. Volcano ages in the Erebus Volcanic Province ...... 80

Figure 4. Crustal thickness for the McMurdo Sound and surrounding region...... 81

Figure 5. Heat flow values for the McMurdo Sound region ...... 82

Figure 6. Airy-Heiskanen at continental scale ...... 83

Figure 7. Vening-Meinesz's isostasy model ...... 84

Figure 8. Schematic illustrating multibeam sonar survey ...... 85

Figure 9. Schematic illustrating multichannel marine seismic reflection survey ...... 86

Figure 10. Workflow for this study...... 87

Figure 11. Schematic of seismic sequence boundary ...... 88

Figure 12. General reflection configurations ...... 89

Figure 13. Reflection configurations of clinoforms ...... 90

Figure 14. Seafloor multiples from seismic line NBP0401-118 ...... 91

Figure 15. Seismic reflector pull up due to igneous body ...... 92

Figure 16. Four geographic regions defined in McMurdo Sound ...... 93

Figure 17. Examples of seismic facies: F1 to F4 ...... 94

Figure 18. Examples of seismic facies: F5 to F7 ...... 95

Figure 19. Spatial distribution of Seismic Facies F1 ...... 96 xii

Figure 20. Lateral seismic facies transition from F2 to F6 in seismic line IT90a-70...... 97

Figure 21. Spatial distribution of Seismic Facies F2 ...... 98

Figure 22. Spatial distribution of Seismic Facies F3 ...... 99

Figure 23: Spatial distribution of Seismic Facies F4 ...... 100

Figure 24. Spatial distribution of Seismic Facies F5 ...... 101

Figure 25. Spatial distribution of Seismic Facies F6 ...... 102

Figure 26. Spatial distribution of Seismic Facies F7 ...... 103

Figure 27. Alternation of diamictite and diatomite and associated seismic reflection, velocity, and density properties at AND-1B core ...... 104

Figure 28. Interglacial sedimentation model in McMurdo Sound ...... 105

Figure 29. Rj seismic surface depth and isochore map of U1 ...... 106

Figure 30. Rj1 seismic surface depth and isochore map of U2 ...... 107

Figure 31. Rk seismic surface depth and isochore map of U3...... 108

Figure 32. Rk1 seismic surface depth and isochore map of U4...... 109

Figure 33. Rk2 seismic surface depth and isochore map of U5...... 110

Figure 34. Location of seismic lines described in Figures 35 – 46 ...... 111

Figure 35. Seismic line NBP0401-100 ...... 112

Figure 36. Seismic line NBP0401-151d ...... 113

Figure 37. Seismic line IT90a-70...... 114

Figure 38. Seismic line IT90a-74...... 115

Figure 39. Seismic line IT90a-73...... 116

Figure 40. Seismic line NBP0401-128 ...... 117

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Figure 41. Seismic line NBP0401-126 ...... 118

Figure 42. Seismic line NBP0401-159 ...... 119

Figure 43. Seismic line NBP0401-157 ...... 120

Figure 44. Seismic line NBP0401-158 ...... 121

Figure 45. Seismic line NBP0401-130m ...... 122

Figure 46. Seismic line NBP0401-125 ...... 123

Figure 47. Seismic facies distribution within seismic unit U1 ...... 124

Figure 48. Seismic facies distribution within seismic unit U2 ...... 125

Figure 49. Seismic facies distribution within seismic unit U3 ...... 126

Figure 50. Seismic facies distribution within seismic unit U4 ...... 127

Figure 51. Seismic facies distribution within seismic unit U5 ...... 128

Figure 52. Isochore of two seismic sequences ...... 129

Figure 53. Isochore of flexural basin sequence from Rj to seafloor ...... 130

Figure 54. Location of seismic lines described in Figures 55 – 61 ...... 131

Figure 55. Relative age from seismic line NBP0401-151g ...... 132

Figure 56. Relative age from seismic line IT90a-70 ...... 133

Figure 57. Geomorphology and age data show Mt. Erebus collapse calderas ...... 134

Figure 58. FR1 in seismic line NBP0401-126 ...... 135

Figure 59. FR2 in seismic line NBP0401-157 ...... 136

Figure 60. FR3 in seismic line NBP0401-158 ...... 137

Figure 61. FR4 in seismic line NBP0401-123 ...... 138

Figure 62. Age constraints of seismic surfaces around AND-1B drillsite ...... 139

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Figure 63. Jump correlation for seismic line NBP0401-159 and MIS-1 ...... 140

Figure 64. Velocity model ...... 141

Figure 65. Spatial extension of the Mt. Erebus, Mt. Terror, and Mt. Bird loads ...... 142

Figure 66. 3D View of Ross Island volcanoes derived from digital elevation model .... 143

Figure 67. Modelling results showing surface deformation from Ross Island loading for increasing lithospheric Effective Elastic Thickness ...... 144

Figure 68. Modelling result showing surface deformation for Mt. Bird loading for a lithospheric Effective Elastic Thickness of 4 km ...... 145

Figure 69. Modelling result showing surface deformation for Mt. Bird and Mt. Terror loading for a lithospheric Effective Elastic Thickness of 4 km ...... 146

Figure 70. Modelling result showing surface deformation for Mt. Bird, Mt. Terror, and

Mt. Erebus loading for a lithospheric Effective Elastic Thickness of 4 km ...... 147

Figure 71. Best fit model ...... 148

Figure 72. Comparison of Rj surface depth from seismic mapping with flexural surface in response to Mt. Bird loading when Te = 4 km ...... 149

Figure 73. Comparison of Rk surface depth from seismic mapping with flexural surface in response to Mt. Erebus loading when Te = 4 km ...... 150

Figure 74. Interpretation of seismic line NBP0401-128, NBP0401-157, NBP0401-158 showing basin sequences ...... 151

Figure 75. Stratigraphic pattern in response to island chain loading ...... 152

Figure 76. Map view of 3D flexural modeling results from Aitken et al. (2012) ...... 153

Figure 77. Profile view of 3D flexural modeling results from Aitken et al. (2012) ...... 154

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Chapter 1: Introduction

Lithospheric flexure induced by extrusive volcanism plays an important role in regional scale basin evolution. Flexural features can help us understand regional lithospheric strength. 3D forward models have been applied for volcano loading to examine the strength of the lithosphere (Ali and Watts, 2003; Watts and Ten Brink,

1989).

Ross Island volcano loading provides a natural experiment to examine the lithospheric strength of the McMurdo Sound region. Ross Island is located at the south end of the Terror Rift, which is the youngest and westernmost component of the West

Antarctica Rift System and recently active with crustal thinning and extension. Ross

Island mainly consists of three volcanoes formed from late Neogene to Holocene

(Armstrong, 1978; Esser et al., 2004). In the McMurdo Sound region, bathymetry data shows a flexural moat around Ross Island that is several hundred meters deeper than the adjacent seafloor and seismic reflection data reveals older flexural surfaces below the modern seafloor (Horgan et al., 2005).

Previous studies of volcano loading in the McMurdo Sound region used spatially limited bathymetry, seismic and gravity data as model constraints (Stern et al., 1991;

Aitken et al., 2011). Abundant marine geophysical data are now available around the western and northern sides of Ross Island. It is necessary to integrate these data to provide better, regional constraints for flexural modeling.

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The goal of this research is to examine the influence on regional basin evolution from Ross Island volcano loading, and to apply a 3D continuous elastic thin plate model to evaluate regional lithospheric strength. This study uses seismic reflection data and bathymetry data to map flexural surfaces and sediment thickness in the flexural moat basin; builds a model to calculate flexural surfaces based on topographic loading and lithospheric effective elastic thickness (Te); and compares calculated flexure surfaces with flexural surfaces mapped by seismic data to get the best fit Te. Study results provide a stratigraphic architectural view of Neogene basins in McMurdo Sound and estimates regional lithospheric strength.

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Chapter 2: Literature Review

2.1 Geological Setting

2.1.1 The West Antarctic Rift System

The West Antarctic Rift System (WARS) (Figure 1) is a Mesozoic – Cenozoic asymmetric rift system, extending from the continental shelf of the Ross Sea, inland beneath the Ross Ice Shelf and part of the West Antarctic Ice Sheet (Fitzgerald, 2002).

The size of the WARS is similar to the Basin and Range in North America, and the East

African Rift (Behrendt, 1999). The east boundary of the WARS is Marie Byrd Land, which is characterized by semi-continuous volcanism since 30 Ma (Hole and Lemaurier,

1994). The west boundary of the WARS is the Transantarctic Mountains, which are argued to form due to thermal and flexural uplift in an extensional tectonic regime (Stern and Ten Brink, 1989). Rifting has resulted in a set of N - S trending basins separated by two basement highs, namely from west to east, Victoria Land Basin and Northern Basin,

Coulman High, Central Trough, Central High, and Eastern Basin (Behrendt, 1999).

The rifting of WARS has been interpreted to occur in two major pulses after the breakup of Gondwana (Schmidt and Rowley, 1986; Stump and Fitzgerald, 1992; Elliot,

1992; Storey, 1996). The first major pulse, which occurred in the Late Cretaceous, caused extensive crustal thinning in the entire WARS and formed N – S trending rift basins

(Cooper et al., 1987a; Salvini et al., 1997). The second major pulse, mainly confined to

3 the western Ross Sea, is linked with the uplift of the Transantarctic Mountains starting at about 50 Ma (Rossetti et al., 2003; Fitzgerald et al., 2006; Storti et al., 2008).

2.1.2 The Victoria Land Basin and Terror Rift

The Victoria Land Basin (Figure 1) is the westernmost basin within the WARS, bounded by the Transantarctic Mountains on the west side and by Coulman High on the east side. The basin is filled with up to 14 km of sedimentary rocks (Cooper et al.,

1987b). Seismic reflection data and core drilling data are interpreted to record 4 phases of sediment deposition in the southern Victoria Land Basin from Eocene to present

(Fielding et al., 2008). The last phase (the Renewed Rifting phase, Middle Miocene to

Recent), is linked with volcanism of the Erebus Volcanic Province and faulting along the

Terror Rift, formed within the Victoria Land Basin.

The Terror Rift is an N-S trending fault zone extending ~ 275km along strike between Mt. Melbourne and Ross Island. The Terror Rift is ~ 75km wide including

Discovery Graben and Lee Arch (Cooper et al., 1987a). From 7 - 10 km (6 - 8%) of extension was reconstructed since ~13.6 Ma based on balanced cross sections in the

Terror Rift (Magee, 2012).

2.1.3 Ross Island Volcanic Complex

Ross Island (Figure 2), part of the Erebus Volcanic Province (Kyle, 1990), is located at the southern end of Terror Rift. Ross Island consists of four volcanic centers ranging in age from Pliocene to Recent (Figure 3): Mt. Erebus (1.31 Ma – present) (Esser et al., 2004), Hut Point Peninsula (1.34 – 0.44 Ma) (Kyle, 1981), Mt. Terror (1.75 – 0.8

Ma) (Armstrong, 1978), and Mt. Bird (4.62 – 3.08 Ma) (Armstrong, 1978). Mt. Bird, Mt.

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Terror and Hut Point Peninsula are arranged almost radially around Mt. Erebus at about

120°, suggesting a pattern related to doming above a mantle plume (Kyle et al. 1992).

Mt. Erebus, an active polygenetic stratovolcano, mainly consists of anorthoclase- phyric phonolite flows (Armstrong, 1978). The oldest dated sample (1,311 ± 16 Ma) is from a dike at Cape Barne (Esser et al., 2004), and a lava lake is still active in the summit crater. Samples dated at various locations on Mt. Erebus using the 40퐴푟/39퐴푟 method suggest three stages of volcano evolution, including the main ‘proto-Erebus’ shield volcano growth from 1.3-0.75 Ma, the proto-Erebus cone-building phase up to 0.25 Ma, and the modern Erebus cone-building phase up to the present (Esser et al., 2004).

Hut Point Peninsula is an elongated volcanic complex about 20 km long and 2-4 km wide extending southwestward from the southern flank of Mt. Erebus. Rock composition ranges from basanite to phonolite with ages between 0.44 ± 0.1 and 1.34 ±

0.23 Ma (Armstrong, 1978; Kyle, 1981). In this study, Hut Point Peninsula is considered part of Mt. Erebus.

Mt. Terror is a large basaltic shield volcano with parasitic cones dating from 1.75

± 0.3 Ma to 0.82 ± 0.14 Ma (Armstrong, 1978). Mt. Terra Nova (0.8 ± 0.5 Ma)

(Armstrong, 1978) is a large basaltic cone located on the west flank of Mt. Terror and is considered in this study to be part of Mt. Terror.

Mt. Bird is a basaltic shield volcano with basaltic and phonolitic parasitic cones dating from 4.62 ± 0.6 Ma to 3.08 ± 0.15 Ma (Armstrong, 1978).

2.1.4 Regional Lithospheric Strength

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Two important factors that influence the strength of continental lithosphere are the thickness and temperature of the lithosphere (Tesauro et al., 2013). Lithosphere is the strong shell of the Earth, ‘floating’ on weaker asthenosphere. The difference between the lithosphere and the asthenosphere is based on mechanical properties, in contrast to the difference between the crust and the mantle, which is based on chemical composition.

Lithosphere is the most important feature in response to loading on million year time scales. The lithosphere-asthenosphere boundary is difficult to map seismically (Kind et al., 2012) because it represents a temperature transition. Only crustal thickness has been investigated by seismic techniques in the study region, so this is discussed in the next paragraph.

Researchers, using receiver function and surface wave phase velocities derived from teleseismic data, show crustal thickness decreasing from 40 km near the

Transantarctic Mountains to 20 km on Ross Island (Figure 4) (Bannister et al., 2003;

Lawrence et al., 2006). Marine seismic reflection and refraction surveys reveal 17 – 23 km thick crust at various locations near Ross Island (Figure 4) (McGinnis et al., 1985;

Cooper et al., 1997). These results document thin crust, which agrees with the known rift tectonic setting characterized by crustal extension and thinning.

Heat flow in the study area (Figure 5) (Blackman et al., 1987; Morin et al., 2010;

Schroder et al., 2011) is significantly higher than the global continental mean 57 mW/푚2

(Sclater et al., 1980). Thermometry of xenoliths included in volcanic rocks support very high temperature (900 – 100 °C) in the middle and lower crust (Berg et al., 1989).

Seismic travel time tomography identified a low-velocity anomaly in the upper mantle

6 beneath Ross Island, which can be caused by a 200 – 300 °C thermal anomaly (Watson et al., 2006). The lithosphere in the study region is characterized by high temperature.

2.2 Isostasy and Lithospheric Flexure Model

The development of the concept of isostasy and lithospheric flexure is well described in the book ‘Isostasy and Flexure of the Lithosphere’ (Watts, 2001). This section is a brief summary from this book from chapter 1 – 3.

The word ‘isostasy’ is derived from the Greek word ‘iso’ meaning ‘equal’.

Isostasy proscribes that the Earth’s lithosphere and asthenosphere tend to reach gravitational equilibrium. Airy (1855) provided a classic gravitational equilibrium model, in which he compared Earth’s crust lying on the mantle as blocks of timber floating on water. Heiskanen (1931) further developed the model to explain large-scale isostatic phenomena (Figure 6). However, these models ignored the strength of the lithosphere.

Vening-Meinesz (1939) introduced a widely accepted model of isostasy including the strength of the lithosphere. In his model, loading on the lithosphere is simplified as loading on an elastic plate floating (the lithosphere) on inviscid fluid (the asthenosphere).

Instead of local compensation as in Airy’s model, the compensation of surface loading is regional in the Vening-Meinesz model. The amplitude and width of the flexure depends on the regional rigidity of the lithosphere (Figure 7).

The solution of a 2D continuous elastic plate model was given by Walcott (1970).

The solution of a 2D discontinuous elastic plate, with discontinuity beneath the loading center, was given by Turcotte (1979), in order to model major faults cutting through

7 crust. The deformation of a discontinuous elastic plate has shorter wavelength but larger amplitude, in response to the same load (Turcotte, 1979).

2.3 Previous Flexural Studies of Ross Island Loading

Stern et al. (1991) compared a 2D continuous plate model with a 2D discontinuous plate model, deformed in response to loading by the Erebus Province volcanoes. The author simplified the real world 3D scenario into a 2D model by assuming an infinitely long line load on an elastic plate, and the discontinuity of the plate is right below the load. The loading topography is averaged from topography between

Mt. Morning and Mt. Erebus. The distance of the flexural bulge to the loading center and the dip angle of strata in seismic lines were used to constrain the model. The author concluded a discontinuous plate with 24 km elastic thickness is the best fit model.

Ten Brink et al. (1997) built a 3D continuous thin elastic plate model to show the flexure caused by loading by the Erebus Province volcanoes. However, the focus of this paper was on formation of the Transantarctic Mountains, and details of the volcano flexure model were not presented.

Aitken et al. (2012) presented a 3D model to simulate flexural surfaces and gravity field anomalies in the McMurdo Sound region, based on loading by Erebus

Province volcanoes. The authors used a continuous thin elastic plate model to calculate flexural surfaces using a spectral method. Two inputs to their model are a Digital

Elevation Model (DEM) of topography to calculate the volcano loads and a set of assumed effective elastic thickness (Te) from 0.5 km to 25 km, which represents the rigidity of the lithosphere. Two outputs of their model are flexural surfaces and free air

8 gravity anomalies based on the deformation of flexural surfaces. Two constraints for their model are seismic reflection data to find the flexural surface (Horgan et al., 2005) and gravity measurements to get the free air anomaly (Blakemore, 2005) on the

McMurdo Ice Shelf. By comparing misfit between the output of the model and the observed result, they concluded that Te is about 2 to 5 km in the study region. They explained the low Te as a combination of high geothermal temperature gradient and preexisting faults. Aitken et al. (2012) compared their modeling data with data constraints from an area south of Ross Island, and pointed out that the DEM and density are two main uncertainties in their model.

Abundant marine seismic reflection data (Figure 2) are available for the ice-free area west and north of Ross Island. These data provide more basin structure profiles that are widely distributed across the flexural basin along the Ross Island margin. This study integrates these data to provide more constraints for flexural volcano loading models.

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Chapter 3: Data

The primary data sources for this study are Digital Elevation Model (DEM) and seismic reflection profiles. As the goal of this research is to quantify deformation due to flexural loading, both surface and subsurface information are needed to:

 Determine the full volume of the volcano loads from DEM and seismic profiles.

 Map the 3D geometry of surfaces that have deflected due to loading from seismic

reflection profiles.

 Establish the volume and geometry of sedimentary fill in the flexural basins from

seismic reflection profiles, which constrains the geometry and evolution of

flexure through time.

3.1 Digital Elevation Model (DEM)

A Digital Elevation Model that integrates land surface topography and seafloor bathymetry is used in this research to quantify the volume of volcano loads.

3.1.1 Land Surface Elevation of Ross Island

Ross Island topographic data were provided by the U.S. Geological Survey and

Land Information New Zealand with 20 m elevation contours for a 1:50,000 scale topographic map. The DEM grid was made from digital contours with the ‘Topo to

Raster’ tool in ArcGIS 10 (Stutz, 2012).

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3.1.2 Bathymetric Data for the Seafloor Surrounding Ross Island

3.1.2.1 Multibeam Bathymetry Data

High resolution bathymetric data were mainly collected with Kongsberg-Simrad

EM 120 multibeam sonar on board R/V Nathaniel B. Palmer during the NBP0401 research cruise in the western Ross Sea (Wilson et al., 2004). Additional data from later

R/V Palmer cruises and R/V Oden cruises are used in this research to cover additional areas. The Marine Geoscience Data System (http://www.marine-geo.org/) hosts all these multibeam bathymetric data. High resolution bathymetric data covers most of the area west and north of Ross Island in McMurdo Sound.

Multibeam sonar produces sound energy in the shape of a fan to map the seafloor.

EM 120 multibeam sonar (Figure 8) has up to 191 beams forming a 150 degree ‘fan’ in the cross-track direction (L-3 Communications, 2000). As the ship moves forward, the seafloor below the ship track is mapped by the beam ‘fan’ of sound energy. The area covered by the ‘fan’ is proportional to water depth. Multibeam sonar records two-way travel time between the ship and the seafloor for each beam in a ship coordinate system.

With proper sound velocity profiles from the water column, angle and travel distance is calculated from the two-way travel time in the ship coordinate system. If the sound velocity profile is significantly in error, the outer beams will show curl up or curl down where there is a flat seafloor. A bathymetric map is made by converting the ship coordinate system to a geographic coordinate system in latitude, longitude, and depth.

The open source bathymetry data processing package MB-System (Caress and

Chayes, 2014) was used to process NBP0401 multibeam sonar data. Raw data are

11 converted to .mb57 file format using MB-system. Sound velocity profiles based on daily measured eXpendable Bathy Thermograph (XBT) are applied to convert two-way travel time to distance. Filters are applied to remove spikes and isolated beams with two criteria: along-track dip departs significantly from the local bathymetry, and all beams with more than 50% flagged beams are rejected. Additional refraction corrections are applied at different water depths at nearly flat seafloor to determine an optimum sound velocity profile. Then this optimum sound velocity profile is applied to reduce distortions at the outer beams (Wilson et al, 2004).

Further manual editing and 3D visualization was done in the commercial software

QPS Fledermaus. A high resolution bathymetric DEM with 20 meter resolution grid was made in ArcGIS (Stutz, 2012).

3.1.2.2 Additional Bathymetry Data

A compilation of all existing bathymetric soundings, including single and multibeam sonar data, was supplemented by inversion of 1-minute satellite altimetry derived gravity data to reconstruct the regional bathymetry of the entire Ross Sea (Black et al., 2011). The grid spacing for this compilation is 100 m. This low resolution bathymetric DEM is used as secondary data to fill gaps in the high-resolution bathymetric coverage.

3.1.3 Integrated DEM

DEMs from 3 different sources were integrated into one DEM using ArcGIS

(Stutz, 2012). The process requires two steps using the ‘Mosaic to New Raster’ tool in

ArcGIS. This tool assigns the value of the top layer where two layers overlap. The first 12 step is to put high resolution multibeam bathymetric DEM on top of low resolution bathymetric DEM to make a new bathymetric DEM. The second step is to put the land surface DEM on top of the new bathymetric DEM to make the final integrated DEM. The grid spacing for this integrated DEM is 20m.

3.2 Seismic Reflection Data

Seismic reflection data, revealing subsurface structure below the seafloor, provides key information about basin structure. The flexural surfaces revealed in seismic reflection profiles are used as data constraints for the flexure models in this study.

3.2.1 Reflection Seismology

Reflection seismology is widely used by geologists to explore the subsurface structure of Earth. Seismic reflection uses a controlled energy source to send waves into the ground, and uses receivers to record the waves reflected back from layered stratum underground. In multichannel marine seismic reflection, air guns are used as the energy source, and a streamer consisting of a set of hydrophones is used as the receiver array

(Figure 9).

Seismic waves are reflected at boundaries where acoustic impedance changes.

Acoustic impedance is defined as the product of a rock unit’s density and acoustic velocity. The primary and most probable reason for acoustic impedance contrast is lithology change, but other factors such as mineral phase changes and gas-oil-fluid boundaries also cause acoustic impedance changes.

The typical 2D multichannel seismic processing workflow includes three major steps: deconvolution, stacking, and migration (Yilmaz, 2001). With proper processing, 13 seismic waves recorded by instruments can be output as a seismic profile, analogous to a geological cross section. Based on seismic profiles, stratigraphic units and geological structures can be mapped.

3.2.2 Seismic Data Sources

Seismic data used in this study (Figure 2) are mainly from the NBP0401 geophysical survey sponsored by NSF, with additional data from the Italian Antarctic

Program and the USGS. Data were originally loaded into Schlumberger’s Geoframe software by Stuart Henrys at GNS in New Zealand. For the current study, seismic data and previous seismic interpretations for the study region were transferred to

Schlumberger’s Petrel software using the plugin tool ‘Geoframe Data Connector’.

The NBP0401 geophysical survey was conducted from 19 January to 18 February

2004 on R/V Nathanial B Palmer. The information in this paragraph is summarized from the cruise report that described seismic acquisition parameters for the NBP0401 survey

(Wilson et al., 2014). The energy source was a 6 air gun array with 1260 cu total volume towed 30 meters behind the ship. A 1200 meter long 48-channel streamer was deployed with 93.2 meters near-offset and 1268.2 meters far-offset. For most multichannel seismic lines, data are recorded 8 seconds long at 2 milliseconds sampling ratio, while the air gun array shot every 12 seconds at a ship speed of 4.5 knots (24 fold).

Seismic processing parameters for NBP0401 survey are described in detail in

Whittaker (2005); a brief review of processing parameters based on this follows. Seismic data was processed in GLOBE CLARITAS software from the Institute of Geological and

Nuclear Sciences, New Zealand. Initial processing included setting geometry, and 14 applying a band pass filter to remove noise. Pre-stack processing includes inside muting to suppress the near-offset seafloor multiple, f-k filtering to remove coherent noise, scaling of 1 dB/sec to recover amplitude of the lower seismic section, applying deconvolution with a gap length of 20 msec / gate length of 50 msec, and applying normal move out correction with velocity picked from velocity spectrum plots. Post- stack processing includes time-variable band pass filtering to suppress the seafloor multiple, deconvolution with 20 msec interval, migration to recover true geometry and auto gain control to balance amplitude between the seafloor and the remaining section.

3.2.3 Mis-tie Analysis

The seismic lines used in this study are from different surveys. The difference in acquisition instruments, and data processing parameters causes offset of the same reflector where two lines cross. Mis-tie analysis is needed to correct the offsets before seismic interpretation.

The ‘Mis-tie manager’ in Petrel is used to conduct mis-tie analysis in two steps.

First, mis-tie analysis is applied for lines from the same survey by correlating both amplitude and phase, and this step is repeated for all the surveys. Second, with the mis-tie corrected NBP0401 lines as the reference, mis-tie analysis is applied again by correlating amplitude from lines in all the surveys. The reason for fixing the NBP0401 survey as the reference is because most lines in this research are from this survey. The reason for correlating amplitude in the second step of mis-tie analysis, unlike correlating both amplitude and phase in the first step, is that seismic data from different surveys have different phase. 15

Chapter 4: Methods

4.1 Workflow for this Study

The workflow for this research is (Figure 10) divided into two major steps: observation and modeling.

The observation is mainly derived from seismic reflection data, and the methodology for seismic reflection mapping is further divided into three steps. Step 1a is to define seismic facies based on seismic reflection characteristics in seismic profiles in the study region. The result of seismic facies analysis is presented in section 5.1.1. Step

1b is to map seismic stratigraphic units bounded by regional seismic surfaces. Step 1c is to further analyze the seismic stratigraphic units and their bounding surfaces by compiling seismic facies maps for each unit, creating isochore maps for each unit in two- way travel time, mapping the depth of each seismic surface in two-way travel time, and defining the maximum extension of volcanos below the seafloor in two-way travel time.

The results of steps 1b and 1c are presented in section 5.1.2. The technique used in seismic reflection mapping is explained in section 4.2.

The modeling section is divided into three steps. Step 2a is to calculate the volume of volcanos as the main input for the model, using the DEM of volcanoes above the seafloor with the seismically-defined extension of volcanoes below the seafloor. To derive the extension of volcanoes below the seafloor, a velocity model is used to convert two-way travel time to depth. The result of the velocity model is presented in section

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5.2.1. Model input parameters are presented in section 5.2.2. Step 2b is to build the elastic thin plate model to output flexural surfaces. Model output results are presented in section 5.2.3. Step 2c is to compare modeling results with observations, to find the best fit model, presented in section 5.2.4. The mathematical foundation of the elastic modeling method is explained in section 4.3.

4.2 Seismic Interpretation

In section 3.2.1, I introduced the basic principles of how seismic reflection data are collected and processed. The outcome of processed 2D seismic profiles is reflected wiggle traces, not geological cross sections. Interpretation is needed to convert wiggle traces in seismic profiles into depositional units and geological structures.

4.2.1 Seismic Stratigraphic Techniques

Seismic interpretation is a widely used tool to explore subsurface geology, and recommended procedures are presented in various publications. Veeken and Moerkeken

(2013) explained in detail about seismic stratigraphy interpretation procedure originally proposed by Vail (1987); the following is a brief review from their publication that is related to my study.

4.2.1.1 Seismic Units

The first step in seismic interpretation is to define seismic units bounded by unconformities. Unconformities are identified based on reflection termination patterns.

Several types of geometric relationships are shown at unconformable surfaces (Figure 11).

The upper boundary of a seismic unit may show erosional truncation indicative of regional erosion, toplap which marks either non deposition or only minor erosion, or concordance 17 with underlying strata that are parallel to the boundary. Concordance is commonly identified by tracing the surface back to an area where there is angular unconformity. The lower boundary of a seismic sequence may show onlap, where horizontal strata progressively overstep a boundary, downlap where inclined strata terminate against the underlying surface, or concordance where the boundary surface and underlying strata are parallel.

4.2.1.2 Seismic Facies Analysis

Seismic facies are defined by characteristic features of reflectors. The goal of seismic facies analysis is to infer depositional environment and lithology. The most commonly used parameters are reflection configuration, continuity, amplitude (horizontal excursion from the time axis), frequency content (vertical separation between zero- crossing on the same seismic trace), and external geometry.

Reflection configuration is represented by geometry of reflectors in a seismic facies unit. Several types of reflection configurations and their possible geological interpretations were summarized by Mitchum et al. (1977) and Schlaf et al. (2005) and are shown in Figures 12 and 13.

 Parallel to wavy reflection configuration: uniform rate of deposition, relatively

stable tectonic setting.

 Divergent reflection configuration: progressive tilting and subsidence during

sediment deposition, forming wedge shape external geometry.

 Hummocky reflection configuration: mark sediment lobes.

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 Chaotic reflection configuration: interpretations include high-energy deposits,

post-depositional deformation, and volcanic rocks.

 Reflection-free reflection configuration: the unit has weak acoustic impedance

contrast, so can be interpreted as sedimentation in a homogenous depositional

environment, salt, or as volcanic rock.

 Prograding reflection configuration (clinoform): formed through prograding slope

deposit.

In seismic facies analysis, four other parameters also provide valuable information on depositional environment and lithology:

 Continuity gives information on lateral changes in energy level of the depositional

environment.

 Amplitude gives information on acoustic impedance contrast across the reflector.

 Frequency (vertical separation of reflectors, not the frequency from signal

processing point of view) gives information about thickness of the beds, however,

care should be taken due to the potential pitfall from thin bed interference

(Widess, 1973).

 External geometry, or shape, of a seismic facies unit provides can be used to infer

depositional environment.

4.2.2 Artifacts in Seismic Data

Although a number of processes have been applied to suppress the seafloor multiple, it is strong in most of the seismic lines from the study region. On the continental shelf in Antarctica, surficial sediments are typically overconsolidated due to ice sheet 19 grounding during the Last Glacial Maximum. The acoustic impedance contrast between water and well-consolidated surficial sediments at the seafloor is enormously high, such that a large amount of energy is reflected at this interface (Figure 14). The seafloor multiple in seismic lines shows double travel time compared with the primary seafloor reflection, and has opposite polarity. Interpretation through and below the seafloor multiple is less confident than in sections above the seafloor multiple.

Another prevailing artifact is the pull up effect caused by igneous intrusions

(Figure 15). Igneous intrusions are very common, due to widespread Neogene to present active volcanism and the proximity of the study region to the major volcanoes of Ross

Island. Igneous bodies have higher seismic velocity than the surrounding sedimentary rocks. Because of this lateral velocity difference, flat strata appear to curve up where igneous intrusions cross them in seismic time sections.

4.2.3 Seismic Interpretation in Petrel

Several tools in the Petrel software package are used for seismic observations in this study. The ‘Mis-tie manager’ tool is used to correct offsets between different surveys. The ‘Seismic interpretation’ tool is used to interpret seismic horizons. The

‘Make/edit surface’ tool is used to interpolate 2D interpretation to a 3D surface. Based on these interpolated surfaces, the ‘Structural operations --> isochore’ tool is used to make an isochore map between two seismic surfaces. The ‘Velocity model’ tool is used to convert seismic surfaces in two-way travel time to the depth domain.

Loop correlation is applied to check consistency of seismic interpretation wherever data allows. Several seismic lines intersect to form a loop. Starting at seismic 20 line 1, a seismic horizon interpretation is followed between each seismic line in the loop, then back to seismic line 1. If the same horizon is offset at the intersection between two lines, then the interpretation is not consistent.

4.3 Modeling Method

The loading of lithosphere on million-year time scales can be modeled as loading on a thin elastic plate floating on inviscid fluid. Based on the elastic thin plate assumption, the flexural surface can be calculated based on topographic loading by volcanoes.

In this study, the elastic thin plate model is calculated using function ‘gravfft’ in the open-source program GMT (Wessel et al., 2013). Function ‘gravfft’ takes a grid of load topography as input, and outputs flexural surface grids. The resolution, of input and output grids is 20 m in this study.

Function ‘gravfft’ uses the spectral method to calculate flexural surfaces. The mathematical foundation of calculating flexural surfaces with the spectral method is well described in the book ‘Isostasy and Flexure of the Lithosphere’ (Watts, 2001). The following section is a brief summary from this book.

The spectral method regards the lithosphere as a filter, taking high-amplitude, short-wavelength loading as input and creating low-amplitude, long-wavelength flexure as output. Fourier analysis is used to break down an arbitrary loading into spectral components.

퐻푘 = 퐹퐹푇(퐻) (1)

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퐻 is the loading topography grid. 퐻푘 is the loading topography in the wavenumber domain. FFT is forward fast Fourier transform.

4 (휌푙− 휌푤) D × k −1 푌푘 = × 퐻푘 × [ + 1 ] (2-1) (휌푚−휌푖) (휌푚−휌푖)× g

푌푘 is the flexure surface in the wavenumber domain, 휌푙 is the density of loading material. 휌푤 is the density of water/air whose position was replaced by loading material.

휌푚 is the density of the mantle. 휌푖 is the density of infill material in the flexural moat. g is the gravitational acceleration. D is the flexural rigidity.

E × Te3 D = (2-2) 12 ×(1−ν)

E is Young’s Modulus. ν is Poisson’s Ratio. Te is effective elastic thickness, which depends on lithosphere thickness, temperature and preexisting structure.

푌 = 퐼퐹퐹푇(푌푘) (3)

푌푘 is the flexure surface in the wavenumber domain. 푌 is the flexure surface grid output. IFFT is inverse fast Fourier transform.

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Chapter 5: Results

5.1 Seismic Framework

This study aims to build on the Neogene seismic stratigraphy established for the southwestern Ross Sea by Whittaker (2005), Horgan et al. (2005), Henrys et al. (2007), and Fielding et al. (2008) (Table 1), by examining the youngest seismic sequences associated with the flexural basin formed around the volcanoes of Ross Island. The lowermost reflector associated with development of the oldest volcano of Ross Island, Mt

Bird, was identified as the reflector denoted as ‘Rj’ by investigating relations between seismic reflection surfaces and the apron of the shield volcano, which formed during the main phase of volcano building. As discussed in the section below establishing the ages of the seismic sequence, Rj is the youngest seismic surface sitting below the apron of Mt.

Bird, so this study has systematically mapped Rj and overlying reflectors up to the seafloor to establish the geometry and seismic character of the infill of the flexural moat basin. Four geographic regions are defined by modern day seafloor geomorphology from west to east, namely the Western Shelf, the Western Slope, the Erebus Basin, and the

Ross Island Slope (Figure 16). Seismic facies distribution and seismic unit description will refer to these geographic regions.

5.1.1 Seismic Facies

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5.1.1 Description of Seismic Facies

Seven seismic facies were identified in the study area, based on vertical reflector frequency, lateral continuity and amplitude, and on reflector geometry and the form of the bodies they occur in. Table 2 summarizes the characteristics of all seven seismic facies and their possible interpretations. Figure 17 and Figure 18 show examples of each seismic facies in representative seismic lines. Seismic facies 1-4 are facies defined by packages with multiple reflectors, whereas seismic facies 5-7 are facies defined by single layers or bodies with > 50 ms thickness. Figures 19-26 show the spatial distribution of each seismic facies on a basemap showing the modern shelf-slope-basin bathymetry and channel systems mapped by Stutz (2012).

Seismic Facies 1 (F1): F1 has high vertical frequency, high lateral continuity, and high amplitude. F1 has parallel internal reflector geometry and sheet-like external shape.

The occurrence of F1 is high, distributed parallel to the axis of the deepest modern

Erebus basin to the west of Mt. Bird and Mt. Erebus (Figure 19).

Seismic Facies 2 (F2): F2 has low to high vertical frequency, moderate to low lateral continuity, and moderate to low amplitude. F2 internal reflectors are subparallel, with a discontinuous to irregular disrupted form. It is quite common to find F2 transitioning laterally to seismic facies F6 which has low-amplitude, discontinuous reflectors or is nearly transparent (Figure 20). F2 typically occurs in sheet-like bodies.

The F2 facies occurs commonly in 2 regions, to the west of Mt. Erebus and repeatedly to the northwest of Mt. Bird, mainly in the deep part of Erebus Basin (Figure 21).

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Seismic Facies 3 (F3): F3 has high vertical frequency, moderate lateral continuity, and moderate to high amplitude. F3 has wavy internal reflector geometry, with channel- form, lenticular, or mounded shapes. The occurrence of F3 is high, but only in the deep part of the Erebus Basin around Mt Bird volcano (Figure 22).

Seismic Facies (F4): F4 has high vertical frequency, high lateral continuity, and moderate amplitude, with a shingled internal reflector geometry that terminates up and down against bounding surfaces. The occurrence of F4 is low, only located to the north of

Mt. Bird (Figure 23).

Seismic Facies 5 (F5): F5 has low vertical frequency, low lateral continuity, and low to high amplitude. F5 has chaotic internal reflector geometry, typically with hyperbolic point-source diffractions. F5 packages have sheet, mound or tongue shapes.

The occurrence of F5 is high, dominantly distributed on the steep western slopes surrounding Mt. Bird and Mt. Erebus (Figure 24).

Seismic Facies 6 (F6): F6 has low vertical frequency, low lateral continuity, and low amplitude. F6 has either subparallel but irregular very low-amplitude internal reflectors, or is reflection free, and typically forms sheet-shaped bodies. The occurrence of F6 is moderate, in the deep Erebus Basin or along the western slope into the basin, west and northwest of Mt Bird (Figure 25).

Seismic Facies 7 (F7): F7 has low vertical frequency, low lateral continuity, and low amplitude. F7 has chaotic internal reflector geometry within subsurface steep-sided convex-upward bodies, or has antiformal reflector ‘pull-ups’ below surface mounds and fissure ridges on the seafloor. The reflection amplitude at the top of F7 is high. The

25 occurrence of F7 is moderate, with surface mounds distributed along the western slopes of Mt Erebus and Mt Bird and a large subsurface body between Mt Bird and Beaufort

Island (Figure 26).

5.1.1.2 Interpretation of Seismic Facies

Many previous workers have examined seismic facies in the Ross Sea with the aim of understanding the signatures of glaciation (Anderson, 1999; Bartek and Anderson,

1991; De Santis et al., 1995; Brancolini et al., 1995). Piston cores have been widely used to establish the character of sedimentary successions imaged on seismic records, however, cores obtained in the Ross Sea typically sample only the uppermost layers due to the difficulty in penetrating the overcompacted surficial materials. Using core data,

Domack et al. (1999) developed a glaciomarine sedimentation/facies model for the western Ross Sea and Bartek and Anderson (1991) created a model specific to interglacial sedimentation in McMurdo Sound. The ANDRILL project obtained the 1285 m-long AND-1B core on the south side of Ross Island, and the sedimentology and sequence stratigraphy of the Neogene strata in this sequence have been summarized by

Naish et al. (2007), Naish et al. (2009), McKay et al. (2009) and McKay et al. (2012).

Hansaraj et al. (2007) and Hansaraj (2008) looked at the seismic stratigraphy of the region south of Ross Island relative to the AND-1B core. All of these studies provide important background information for interpreting the seismic facies defined in this study.

Seismic stratigraphy and facies analysis of glacial marine deposits reported in the studies cited above show that, in general, ice-proximal deposits have internally chaotic or

26 non-reflective character, commonly occur in wedge-shaped or lenticular bodies, are bounded by erosional unconformities and are commonly associated with clinoforms defining ‘grounding-zone wedges’. More ice-distal glacial marine deposits have stratified character with subparallel high- to medium-amplitude reflectors.

Hansaraj (2008) showed that the lithologies in the AND-1B core that correlate with subparallel, high- to moderate-amplitude reflectors consist of alternating diamictites, deposited subglacially, and diatomites deposited in open-marine conditions (Figure 27) in the Pliocene section of core and of alternating diamictites, mudstones and volcaniclastic material in the Pleistocene part of the core.

Bartek and Anderson (1991) used an array of sediment cores from 0.6 – 2.6 m long to reconstruct sedimentary facies in McMurdo Sound deposited since the retreat of grounded ice from the area approximately 7,000 years BP (Figure 28). Their model shows resedimentation of volcanic material by sediment gravity flows, including turbidity currents, down the steep slope of Ross Island into the deep Erebus Basin.

Pelagic and hemipelagic diatomaceous ooze and mud form a large component of deep

Erebus Basin deposits, with contributions of sand and mud into the basin from the west via channels that incise the shelf and slope. Their results also demonstrated that the

McMurdo Sound region to the west and north of Mt Bird received a component of quart- rich sediment gravity flow material contributed to the basin from the MacKay Glacier, which transects the Transantarctic Mountains.

Seismic facies F1 is strongly laminated, with parallel and continuous high- amplitude reflectors in bodies that are subhorizontal and form sheets. In the glacial

27 marine setting of the Norwegian margin, laminated seismic facies are interpreted as hemipelagic deposits (Nygard et al., 2005). Seismic studies from the Ross Sea showed laminated seismic facies formed in a relatively ice-distal glacial marine setting, representing pulses of coarser-grained glacially-derived sediments alternating with fine- grained biogenic and hemipelagic material (De Santis et al., 1995; Bartek et al., 1997;

Domack et al., 1999). Seismic lines crossing the AND-1B core site show that alternation between open marine diatomite and subglacial diamictite, with components of volcaniclastic material, produces high-amplitude laminated seismic facies (Hansaraj,

2008). Hence, possible interpretations for F1 include: alternation between open marine diatomite/hemipelagic mud and glacially-derived siliciclastic sediments and alternation between open marine diatomite/mud and subglacial diamictite. The localization of F1 facies in the deep part of the Erebus Basin and the sheet-like form of these units indicates a low-energy environment.

Seismic Facies F2 and F6 grade laterally into each other, changing from subparallel, discontinuous moderate-amplitude reflectors to semi-transparent internal geometry. Overall both facies form sheet-like bodies within layered seismic sequences.

Seismic facies with strongly discontinuous to chaotic and reflection-free characteristics have been interpreted to represent glacially-derived diamictons, formed as subglacial tills or ice-proximal ‘till deltas’ (De Santis et al., 1995; Bart and Anderson, 1995; Bartek et al., 1997; Batchelor et al., 2013). Reflection-free intervals can also represent homogeneous deposition of fine-grained material (muds or biogenic oozes) for long time intervals and environments where winnowing by bottom currents results in a more

28 homogenous grain size by removing fines (Bartek et al., 1997). Hansaraj (2008) showed that reflection-free units traceable around the seismic grid south of Ross Island correlated with thick diatomite units in the AND-1B core. The F2 seismic facies is not distributed uniformly throughout the area. One occurrence is on the western slope to the northwest of Mt Bird, where it could represent glaciomarine deposits rich in ice-rafted debris from the west. The other occurrence is in the axial part of the Erebus Basin, in the same region as facies F6, indicating a relatively low-energy environment. These deposits may be distal facies of volcanic mass flows from Ross Island, or may be homogenous materials such as the diatomites in the AND-1B core.

Both the lenticular to mounded internal reflector geometry, and the nested channel-form shapes within the F3 facies point to deposition by channelized flow. The

F3 facies occurs in the deep part of Erebus Basin near Mt. Bird, where piston cores show sediment gravity flow deposits, including turbidites, occur together with biogenic oozes and hemipelagic muds in the interglacial deposits found immediately below the modem day seafloor (Bartek and Anderson, 1991). The F3 facies also coincides with the modern-day axial channel system of the Erebus Basin and occurs where the submarine

Wilson Sea Valley enters the basin. The F3 facies is therefore interpreted to represent sediment gravity flow deposits, probably both volcaniclastic material derived from Ross

Island and quartzose material derived from the Transantarctic Mountains. The mounded shapes could also indicate transport and reworking by bottom currents. The high amplitude reflectors in this facies probably indicate alternation of these types of deposits with fine-grained pelagic/hemipelagic deposits.

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Seismic facies F4 has dipping, parallel reflectors resembling clinoforms.

Clinoforms are characteristic of prograding deposits, common in settings including deltas, submarine fans and glaciomarine grounding-zone wedges. Because the occurrence of F4 is very limited, it is not possible to establish whether the dipping reflectors represent clinoforms or whether some local tilting due to magmatic intrusions and/or faulting could have produced the observed dips. The significance of facies F4 thus can’t be established.

Seismic facies F5 has chaotic internal reflectors, which are commonly attributed to debris flows and also to subglacial tills. The strong correlation between the spatial distribution of F5 with the steep flanks of the Mt Bird and Mt Erebus volcanoes argues that facies F5 proximal to Ross Island represents volcanic debris flow deposits forming part of the submarine apron of the volcanoes. Some F5 bodies clearly thin away from the volcano flank. Chaotic seismic units with hyperbolic diffractions are typical of debris flows containing large blocks within them (e.g. Rees et al., 1993; Nygard et al., 2005).

Where facies F5 occurs at considerable distance from Ross Island, on the western slope northwest of Mt. Bird, it is more probable that they represent inhomogeneous deposits such as diamictites, formed subglacially or in an ice-proximal setting. Facies F5 is interpreted as either volcanic debris flows derived from Ross Island or ice-proximal glaciomarine deposits.

Many of the occurrences of seismic facies F7 mark rounded flat-topped seamounts or linear elongate ridges interpreted as volcanic edifices (Stutz, 2012). The high reflection amplitude at the top of F7 occurrences is probably due to the strong

30 acoustic property contrast between volcanic material and interlayered or surrounding sediments. Large subsurface bodies that are convex-up in form and internally reflection free have a typical shape of subvolcanic intrusions and they occur in spatial proximity to volcanic islands and seamounts. Hence, F7 is interpreted as volcanic subsurface bodies and surface seamounts and fissure ridges.

5.1.2 Description of Seismic Surfaces and Units

Five regional seismic surfaces were identified in the flexural basin for this study, namely Rj, Rj1, Rk, Rk1, and Rk2. Rj and Rk are correlated with seismic surfaces from previous studies (Horgan et al., 2005; Fielding et al., 2008), whereas Rj1, Rk1, and Rk2 are seismic surfaces defined in this study. Five seismic units were defined, bounded by these seismic surfaces, namely, U1, U2, U3, U4, U5. Table 1 lists the seismic surfaces and seismic units defined in this study and their correlations with previous studies. For each seismic surface, the general shape, the relations with the unit below it, the spatial distribution, and the dipping pattern is described in the following paragraphs. The slope defined by surface maps and the apparent dip in seismic lines are used to describe the dipping pattern. The locations of seismic lines used in section 5.1.2 are shown on Figure

34. Seismic facies within each seismic unit are described.

5.1.2.1 Seismic Surfaces

Reflector Rj: Rj is a smooth to wavy surface, generally conformable with the unit below it. The Rj surface deepens towards the north to a maximum depth around 1750 ms and shallows towards the west where it is eroded at the seafloor or truncated by the overlying Rj/Unit 2 or, to the south, beneath Rk/Unit 3 (Figure 29a). Reflector Rj is not 31 present to the southwest of Mt Erebus, but appears to the south of Mt Erebus in the

Windless Bight region.

Two patterns are defined by the Rj surface. Rj slopes toward Mt Bird in the region northeast of Mt. Bird (Figure 29a, Figure 35, Figure 36) and in the region to the west and southwest of Mt Bird (Figure 26a, Figure 40). To the northwest of Mt Bird, however, Rj deepens within a blocky, north-northwest trending zone and the surface slopes into this zone from the west and the east. Overall, the dip and depth pattern of Rj fits the hypothesis of deformation caused by flexural loading, with radially inward slopes toward the volcano center, except the area northwest of Mt. Bird. As discussed below, the deep northwest-trending zone is spatially associated with Terror Rift faulting.

Reflector Rj1: Rj1 is a broadly undulating surface, with local incisions into the underlying seismic unit. The Rj1 surface is less extensive in area than Rj, being eroded at the seafloor to the northwest of Mt Bird, and truncated against the overlying Rk/Unit 3, thinning until it disappears southward. Rj1 has both conformable and angular relations with internal reflectors of the underlying U1 seismic unit in the north, but consistently shows an angular truncation of U1 towards the south.

Rj1 deepens toward the north with a maximum depth of around 1700 ms, and shallows toward the west (Figure 30a). The configuration of the Rj1 surface is very similar to Rj. Rj1 slopes inward toward Mt Bird from the northeast (Figure 35, Figure

36) and from the west (Figure 40), but slopes toward a northwest-trending deep zone to the northwest of Mt Bird.

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Reflector Rk: Rk is a rough undulating surface accompanied with sharp incisions into the underlying unit with depth of 17-43 ms in TWT, or about 17-43 m in height with average velocity of 2000 m/s in sediments, and with width of 0.8-2.6 km. The Rk surface extends throughout the region mapped and has the same erosional character everywhere.

Rk is eroded at the seafloor to the northwest of Mt Bird, from this area southward it is truncated against the overlying Rk1 reflector and overlapped by Unit 4 above Rk1.

Reflector Rk has a lower dip angle and truncates underlying reflectors and seismic units with angular unconformity.

Rk has two different regions where the surface deepens (Figure 31a). One is located to the north-northwest of Mt Bird, with maximum depth around 1400 ms, in the same north-northwest trending angular zone where deepening of Rj and Rj1 reflectors also occurs. Rk dips and deepens eastward along the western flank of Mt Erebus (Figure

40, Figure 43). In contrast, along the west side of Mt Bird, Rk is subhorizontal with a gently concave-upward form (Figure 40).

Reflector Rk1: Rk1 is a relatively smooth surface, that extends regionally throughout the study area. Rk1 and overlaying strata are conformable with the underlying seismic unit and show onlap both to the west and to the east. There is a hint of a deepening of the Rk1 surface to the north of Mt Bird, with a maximum depth around

1300 ms, but the blocky northwest-trending ‘deep zone’ seen in underlying surfaces has largely disappeared in the Rk1 horizon. Over most of the region, Rk1 has a subhorizontal, gently concave-upward shape (Figure 32a). This form is typical to the north and west of Mt Bird and also along the northwest margin of Mt Erebus (Figures

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32a, Figure 40, and Figure 44). The geometry of Rk1 changes along the southwest flank of Mt Erebus, where the surface slopes and deepens toward the southeast (Figure 46).

Reflector Rk2: Rk2 is a broadly undulating surface limited in spatial extent to the west and southwest of Mt. Erebus. Rk2 is subhorizontal and broadly concave upwards.

The surface truncates internal reflectors in the underlying seismic unit.

5.1.1.2 Seismic Units

Seismic Unit U1: U1 is distributed around Mt. Bird, truncated to the west of

Erebus Basin at the seafloor and pinching out against overlying units to the south (Figure

29b). U1 continues northward beyond the study region, but was not mapped there due to data limits and the interests of this study. The bounding surfaces and internal reflectors of U1 transitions eastward into the volcanic apron of Ross Island and can’t be traced beyond this limit. U1 is thickest in a semicircular zone around Mt. Bird, with a maximum thickness at about 275 ms northeast of Mt. Bird. The thickness pattern is irregular with thicker regions extending away from Mt Bird toward the northwest. U1 onlaps older units to the north and west and has a wedge shape thickening toward Mt Bird (Figure 38,

Figure 43, Figure 45). Where U1 is layered, there is a divergent pattern of reflectors with decreasing dip upward in the eastern, thickest part of the wedge (Figure 45, Figure 46).

A generalized map of the seismic facies within U1 is shown in Figure 47.

Seismic facies F5 is distributed continuously along the slope surrounding Mt. Bird, consistent with mass flows sourced from the building of Mt. Bird volcano. Outboard of this zone, a large area of facies F6 is present. These relatively reflector-free deposits could be more distal volcaniclastic mass flow deposits, or could be homogeneous

34 pelagic/hemipelagic material, such as the thick diatomites found in the AND-1B core.

Facies F1 occurs further from the volcano flank to the northwest of Mt Bird, and likely consists of alternating fine-grained hemipelagic/biogenic materials and either mass flows, sourced from Mt Bird or from the Transantarctic Mountains via the Wilson Sea Valley, or glacial diamictites deposited during ice advance cycles, as in the AND-1B core.

Smaller patches of facies F2 and F3 northwest of Mt Bird may also be sediment gravity flow deposits, possibly remolded by bottom currents. Overall, U1 is dominated by volcanogenic gravity flow deposits from Mt. Bird, with likely components of glaciogenic deposits, siliciclastic sediment from the Western Slope and pelagic/hemipelagic layers.

Seismic Unit U2: U2 is distributed around Mt. Bird, with a more limited extent than U1. U2 is eroded at the seafloor to the northwest of Mt Bird and disappears toward the south where it is truncated by the Rk reflector and overlying U3 seismic unit (Figure

30b). To the east U2 merges with the volcanic apron. Like U1, U2 thickens inward toward Mt Bird and has a thick zone extending toward the northwest, where it reaches a maximum thickness of about 425 ms. U2 forms a wedge-shaped body, dipping toward the Bird volcano, with a divergent reflector pattern in the thick part of the wedge and onlap to the west onto the Western Slope (Figure 39, Figure 40, and Figure 41).

The generalized map of seismic facies distribution within U2 (Figure 48) shows a more patchy distribution of seismic facies F5 proximal to the steep slope of Mt Bird, probably consisting of volcanogenic mass flow deposits. These are less abundant than in the underlying U1 strata. Facies F5 also occurs in patches that are far to the northwest of the flank of Mt Bird, and these chaotic deposits are most likely proximal glaciogenic

35 deposits. Seismic facies F3 is distributed around Mt. Bird (Figure 48), likely formed by distal sediment gravity flows deposited by channelized turbidite flows and/or reworked by bottom currents. Several patches of F2, to the northwest of Mt. Bird, may record alternating ice-proximal and ice-distal glaciomarine deposits. A patch of F6 is distributed northwest of Mt. Bird, which is possibly thick diatomite or mud deposited in an open marine setting. A patch of F4 occurs northeast of Mt. Bird, the only place it is present. U2 is the thickest seismic unit and contains the most types of seismic facies.

Seismic Unit U3: U3 is present throughout the Erebus Basin, extending further southward than the underlying units. The western limit of U3 occurs where it is truncated by erosion at the seafloor in the northwest and where it is truncated by Rk and overlapped by seismic unit U4 along most of the western basin margin (Figure 31b). U3 is thin in the basin adjacent to Mt Bird, but shows a northwest-trending thick zone to the northwest of

Mt Bird where the maximum thickness is about 175 ms, a pattern similar to the underlying seismic units. In contrast to the lower units, seismic unit U3 thickens towards

Mt. Erebus. Seismic lines offshore of Mt Bird show that U3 consists of subhorizontal strata that onlap both the Western Slope and the eastern flank of Mt Bird (Figure 40, and

Figure 41). The geometry in the south is different, with U3 forming a wedge-shaped package that is thin and onlapping to the west and thickens to the east toward Mt Erebus

(Figures 42, Figure 43, and Figure 44). Offshore of the area where Mt Bird meets Mt

Erebus, reflectors within U3 are horizontal (Figure 42), whereas further south offshore of

Mt Erebus the reflectors form a divergent pattern dipping toward the volcano.

36

The generalized map of seismic facies distribution within U3 (Figure 49) shows that chaotic facies occur mainly along the Ross Island slope near Mt Erebus, with isolated patches futher north near Mt Bird. The F5 facies probably records the active building of proto-Erebus and the waning of mass wasting from Mt Bird volcano. Seismic facies F2 occurs in the deep basin adjacent to Mt Erebus, but also distal to Mt Bird in the northwest. Both of these regions may consist of relatively ice-proximal glaciomarine deposits, or the facies near Mt Erebus may consist of more volcanogenic sediment gravity flow deposits. Patches of sediment with seismic facies F3 are distributed in the axial part of the basin around Mt. Bird (Figure 49), possibly indicating deposition from sediment gravity flows carrying finer-grained materials, transported along channels funneling material along the basin axis or from the shelf/slope to the west.

Seismic Unit U4: U4 is regionally extensive and generally thinner than the underlying units. U4 is present mostly in the Erebus Basin. U4 pinches out towards the west (Figure 32b) and shows a similar, but less pronounced, thickening pattern toward Mt

Erebus. The zone of thicker strata northwest of Mt Bird seen in all underlying units is not defined in U4. In most of the basin, U4 forms a subhorizontal ‘ponded’ unit that thins and onlaps both to the west onto the Western Slope and to the east toward Ross Island

(Figure 40, Figure 41, Figure 42, and Figure 43). Only in the southern basin, offshore of

Mt Erebus, the unit is thin and onlapping to the west and forms a wedge-shaped body thickening eastward toward Mt. Erebus with a maximum TWT of about 90 ms. In this zone, internal reflectors dip to the east and probably show a divergent pattern, although this is partly obscured by the overlying U5 unit.

37

Seismic unit U4 shows a different seismic facies pattern than all underlying units.

U4 is dominanted by the laminated facies F1 and F3 (Figure 50). Small patches of F5 are distributed west of Mt. Erebus, indicating ongoing volcanogenic mass flows from Mt.

Erebus. Another small patch of F5 is found northeast of Mt. Bird, and is probably derived from a seamount northeast of Mt. Bird. The strong, continuous subhorizontal lamination of F1 is likely formed by alternation of fine-grained biogenic and hemipelagic deposits and subglacial or ice-proximal deposits, including a significant volcanogenic component, as found in the upper section of the AND-1B core. Although U4 is too thin to clearly reveal seismic characteristics around Mt. Bird, seafloor geomorphology shows the axial basin channel system extends through this area, so F3 may represent sediment gravity flow deposits in this area.

Seismic Unit U5: U5 occurs only to the southwest of Mt. Erebus (Figure 33b).

U5 pinches out towards the north and west and extends southeastward where the seismic data has artifacts due to complex volcanic mounds on the seafloor. The isochore map for

U5 was cropped to remove the region with artifacts. The thickness map of U5 shows high lateral variation, due to the presence of the incised channel system and volcanic ridges on the seafloor. The thickest part of U5 in the zone that could be mapped is about 126 ms, excluding volcanic fissure ridges.

U5 consists entirely of the chaotic seismic facies F5 (Figure 51). Since U5 is thicker toward the southeast, F5 is probably a volcanogenic mass flow from the southwest flank of Mt. Erebus.

5.1.2.3 Seismic Sequences

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Based on the similarities in reflector surface geometry and isochore patterns, 2 seismic sequences have been defined. Seismic sequence 1 (SS1) includes seismic units

U1 and U2, bounded by the Rj reflector at the base and the Rk reflector at the top.

Seismic sequence 2 (SS2) includes seismic units U3, U4 and U5, bounded by the Rk reflector at the base and the seafloor at the top. Isochore maps for the 2 sequences are shown in Figure 49a and Figure 49b. An isochore map for the entire stratigraphic sequence related to development of the Bird-Erebus flexural basin is shown in Figure 50.

The isochore map of SS1 clearly shows increased thickness in a zone of approximately circular shape around Mt Bird volcano. SS1 also has a thick zone with a linear, well-defined western boundary extending to the northwest from Mt Bird. SS1 disappears to the south, away from Mt Bird.

The isochore map for SS2 shows a zone of thickening toward Mt Erebus volcano.

SS2 has a thicker zone of more irregular and smaller spatial extent associated with Terror

Rift northwest of Mt Bird.

5.1.3 Age Constraints for Seismic Surfaces and Units

5.1.3.1 Volcano Ages

Previous workers have used isotopic ages from volcanic rocks exposed above sea level in the Erebus Volcanic Province to estimate the ages when volcanoes developed, as summarized in Section 2.1.3 (Figure 3). Most volcanoes of the Erebus Volcanic Province have relatively poor age constraints, due to ice cover and lack of dissection of the volcanoes. Surface flows on Mt Bird are interpreted to constrain the main phase of shield volcano growth between 3.8 – 4.6 Ma, based on dates from cliff exposures near the 39 onshore base of the volcano. Mt Erebus has more abundant dates, interpreted to record building of the terrestrial shield of the volcano beginning ~ 1.3 Ma; the subaqeous portion would, therefore, be older. The main mass of the volcano, called ‘proto-Erebus’ was formed between ~1.3 and 0.75 Ma and the upper part, called ‘modern Erebus’ formed over the last 0.25 Ma (Esser et al., 2004; Harpel et al., 2004). Further age constraints from the volcanic record in offshore seismic profiles are examined here.

5.1.3.2 Age Constraints from Seismic Data

The relative ages of seismic surfaces and volcanic units can be identified by cross-cutting and onlapping relations on seismic profiles. Once the relative age is determined, the isotopic dates from the terrestrial outcrops of the volcanoes can be used to constrain the ages of seismic reflectors and the units they bound. Two types of volcanic bodies are used here to determine relative age: the subsurface ‘apron’ of the volcanic edifices, and fissure ridges on the seafloor that form satellite bodies around the volcanoes. The volcano aprons imaged in the subsurface formed during the main phase of volcano building, whereas the fissure ridges may have formed during any stage of volcano building but, because they form well-defined edifices on the seafloor and are not fully blanketed by sedimentary deposits, they likely formed at a relatively late stage.

5.1.3.3 Age Relations from Volcanic Aprons:

There are clear relations between the subsurface apron of Mt Bird volcano and seismic reflectors. Rj is the youngest surface below the apron of Mt. Bird (Figure 55).

Although signal-to-noise ratio is low below the apron of Mt. Bird, Rj can be clearly traced from a good-quality, high signal-to-noise ratio area into the apron. The age of Rj is

40 therefore assigned to be older than the main phase of Mt. Bird, which is 3.8-4.6 Ma based on dates from volcanic flows in cliffs along the west coast of the volcano (Armstrong,

1978).

The Rj1 reflector can be traced above the volcanic apron of Mt. Bird, and strata in

U2 above Rj1 onlap the top of the apron (Figure 55). In the vicinity of Mt Bird, Rj1 is typically a boundary between chaotic units interpreted to be volcanogenic mass flow deposits below and more laminated sedimentary units above (Figure 56). The age of Rj1 is assigned as younger than the main phase of Mt. Bird, which is 3.8-4.6 Ma.

In the seismic unit U3 between Rk and Rk1, a chaotic unit with seismic facies F5 is present around the northwest flank of Mt. Erebus (Figure 49), which is interpreted to be volcanic debris flow deposits from Mt. Erebus. Because of signal-to-noise ratio issues and artifacts caused by volcanic material on the seafloor proximal to Mt Erebus, no clear relations between seismic reflectors and the Erebus volcanic apron could be reliably established elsewhere.

An age constraint for the seismic unit U5, above reflector Rk2, can be inferred from the dated volcanic history of Mt Erebus. U5, which has chaotic internal character and forms distinctive hummocky terrain on the seafloor, is interpreted to be a volcanogenic mass flow derived from the southwest flank of Mt Erebus. Interpretations of the geomorphology and age dates on lavas from the upper sector of Erebus suggest a massive sector or caldera collapse of proto-Erebus volcano between 750,000 - 700,000 years ago, and 2 probable caldera collapses from ~90,000 – 11,000 years ago (Figure 57)

(Esser et al., 2004; Harpel et al., 2004). Because U5 is the uppermost unit immediately

41 below the seafloor, it is unlikely to be the 750,000 collapse event which, based on the

AND-1B core, would be covered by >100 m of strata. Given the geometry of the missing caldera rims, it seems most likely that U5 was part of the debris derived from the failed caldera at ~90,000-70,000 years ago.

5.1.3.4 Age Relations from Fissure Ridges:

Four volcanic bodies that form fissure ridges (FR) along the western margin of

Ross Island are used to show examples of age relations with the seismic stratigraphic units. FR1 and FR2 are aligned parallel to a radius to the center of Mt. Erebus and, therefore, are probably formed at the same time as the volcano, whereas FR3 and FR4 have N-S trends, unrelated to Erebus geometry

FR1 (Figure 58) and FR4 (Figure 61) are examples where both Rk and Rk1 can be traced below the surface ridges, indicating FR1 and FR4 are younger than both reflectors. FR2 (Figure 59) and FR3 (Figure 60) are both onlapped by the Rk1 surface, whereas the Rk surface passes below them, therefore these fissure ridges as younger than

Rk and older than Rk1. These relations show that all fissure ridges formed younger than the age of Rk, hence, the age of Rk is assigned as older than Mt. Erebus, which is

1.31Ma-present (Esser et al., 2004).

5.1.3.5 Age Constraints from Core

The ANDRILL program acquired the AND-1B core through strata inferred to be within the flexural basin to the south of Mt. Erebus. Chronology data from AND-1B core

(Wilson et al., 2012) (Figure 62) provides age constraints for seismic surfaces mapped through a local seismic grid around the drillsite. The seismic data on the west side of

42

Ross Island mapped in this study can’t be directly connected with the seismic grid because the Hut Point Peninsula extension of Mt Erebus separates them. In order to use age constraints from AND-1B core, a ‘jump correlation’ between seismic surfaces west of Ross Island and south of Ross Island is required.

Here we assume that the stratigraphic section in the deepest part of the flexural basin west of Ross Island correlates with the section at the AND-1B drill site south of

Ross Island, located in the same axial basin position. Seismic line 157 west of Ross

Island is used to correlate with seismic line MIS_1 south of Ross Island (Figure 63).

Seismic line 157 from west of Ross Island is shifted down about 60 ms in the vertical axis such that depth to the seafloor matches. This results in a good depth match of Rk and

Rj between the two lines. The thicknesses between Rk and the seafloor range from 180-

150 ms in 157 and 180-130 ms in MIS_1. The thicknesses between Rj and Rk range from

150-80 ms in 157 and 145-120 ms in MIS_1. These unit thicknesses also are a reasonable match.

In the area south of Mt. Erebus, Rj is a smooth surface, and Rk is a rough erosional surface, the same characteristics for these two surfaces to the west of Ross

Island. According to the seismic facies classification defined in this study, the seismic facies between Rj and Rk is the laminated facies F1, and between Rk and seafloor is facies F2 on the MIS_1 line. To the west of Ross Island, units U1 and U2 are missing to the southwest of Ross Island. Seismic units U3 and U4 consist of seismic facies F2 and

F1, comparable to those above the Rk surface to the south of Ross Island.

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The depth and thickness of units, the character of the major reflectors, and the seismic unit characteristics support the correlation of Rj and Rk surfaces west and south of Ross Island. If this ‘jump correlation’ is correct, it indicates that strata between Rj and

Rk range from ~3.2 – 2.0 Ma, and strata from Rk to the seafloor range from 2.0 to <0.25

Ma (Wilson et al, 2012; McKay et al., 2012). The age from the core for Rj of ~3.2 Ma is younger than the >3.8-4.6 age implied by applying the age dates for terrestrial volcanic rocks to the subsurface volcanic apron imaged seismically. The age for Rk of ~2.0 Ma is older than the >1.3 Ma age determine from the age of dated terrestrial samples for Mt

Erebus.

5.2 Modelling

A 3D thin elastic plate model was built to simulate flexural deformation caused by loading from the volcanoes of Ross Island. A velocity model was developed to convert the mapped seismic surface Rj, at the base of the flexural basin, from two-way travel time to depth. The depth to Rj is used in both the model input and in the analysis of the model results.

5.2.1 Velocity Model

The seismic reflection method records two-way travel time (TWT), so by default

TWT is the vertical axis for seismic images. For this research, the seismic surface Rj in seismic profiles needs to be converted from TWT into depth. The maximum depth of Rj is needed to calculate the volume of volcano loads. The dip angle of Rj is needed for evaluation of model results.

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Two data sets from the AND-1B drilling project are available to develop a velocity model for McMurdo Sound area seismic profiles. A Vertical Seismic Profile

(VSP) survey was conducted at the AND-1B borehole (Hansaraj, 2008). P-wave velocity was measured on the whole core scan for the 1285 m-long AND-1B core, and yielded a linear trend from 1791 m/s to 3188 m/s from the top to the bottom of the core (Niessen et al, 2007). A time/depth curve from the core P-wave velocity measurements is compared with the first arrival times from the VSP experiment and the two curves match very well

(Figure 64).

For this research, a two-layer velocity model is constructed. The first layer is the water column with constant velocity of 1500 m/s. The second layer is a linear velocity model for sedimentary strata of: V = 1791 + 0.92 × Z , where Z is sediment thickness. The time/depth curve from this velocity model is plotted with VSP and whole core scan data (Figure 64) and the model curve fits well with these results.

The AND-1B core is located in the flexural moat to the south of Ross Island.

Since most of the area mapped in this study is also in the flexural moat from loading of

Ross Island, though to the west and north of Ross Island, it is reasonable to apply to time/depth conversion from the AND-1B drillsite to the seismic data from the study area.

5.2.2 Model Input

The volume of volcano loads calculated from the DEM is the key input for the flexure model. The spatial extent of each volcano load is defined at the edge of the volcanic apron as observed on bathymetric and seismic data (Figure 65, and Figure 66), and the elevation of each volcano is extracted from the DEM. In order to calculate 45 volume, the base of the volcano needs to be defined. Substantial parts of the volcano aprons are buried by sediments, so the base of the volcanoes is not the seafloor. The Rj seismic surface is identified in this study as the youngest surface before formation of

Ross Island volcanoes, so the Rj surface is used as the base of Ross Island volcano loads.

The deepest part of Rj is mapped on the west of Ross Island from seismic line NBP0401-

157 at 1540.41ms and south of Ross Island, on seismic line MIS_1 at 1489.67ms.

Applying the velocity model discussed in the last section to convert time to depth, the average depth of Rj surface is at -1200.12 m. Hence, -1200m is used as the base of the

Ross Island volcanoes.

Based on the dimensional parameters above, the volume of Mt. Bird is 9.256 ×

1011 푚3, the volume of Mt. Terror is 3.238 × 1012 푚3, the volume of Mt. Erebus is

2.960 × 1012 푚3, and the total volume of Ross Island is 7.124 × 1012 푚3 (Table 3).

Table 3 also shows the volume of Ross Island volcanos calculated by Esser et al. (2004), which is systematically smaller than the volume calculated in this study. The main reason is they use -500 m from average seafloor depth as the base of the volcanoes, whereas I use -1200 m from seismic mapping as the base of the volcanoes.

The parameters used in the model are listed in Table 4. Effective Elastic

Thickness (Te) is the key parameter to be tested by the modeling, to constrain the strength of the lithosphere. Te is tested by using a set of values from 1 to 40 km with 1

3 km increments. The density of loading material 휌푙 is 2650 kg/푚 , which is from a gravity model near Hut Point Peninsula (Melhuish et al., 1995). The flexural moat is largely filled by water instead of sediments, hence, the density of infill material 휌푖 is set

46 to the density of ocean water. The density of mantle 휌푚 is set to the relatively low value of 3260 kg/푚3, due to hot mantle in this region (Aitken et al., 2012). Young’s Modulus E is 1011 Pa, and Poisson’s Ratio is 0.25.

5.2.3 Model Output

The modeling produced a total 120 outputs, results representing 1 to 40 km Te for

3 volcano loads. Based on these 120 outputs, 40 flexural surfaces responding to total

Ross Island loading are produced by adding the 3 flexural surfaces for each load.

Figure 67 shows flexural surfaces from Ross Island loading with 1km, 5km,

10km, and 25km Te. The flexural surfaces derived from weaker lithosphere with smaller

Te shows deeper and narrower flexure or, in other words, higher amplitude and lower wavelength. Results for flexural surfaces modeled using stronger lithosphere with larger

Te show shallower and wider basins or, in other words, lower amplitude and higher wavelength. When Te is 1km, the flexural surface shows three separate sub-basins below each volcano. The maximum depth is -2400m for the sub-basin below Mt. Erebus. The flexural bulge occurs between the two zero contour lines on the maps. When Te is 5km, the flexural surface summed for all Ross Island volcano loads shows a single east-west elongated basin with maximum depth of -1200 m located in the area between Mt. Erebus and Mt. Terror, with the flexural bulge located outside of the map region of this study.

When Te is more than 10 km, the total-load flexural surface shows a single basin with a circular pattern, with the flexural bulge outside of the map region of this study. The maximum depth is -600m for Te = 10km, and -200m for Te = 25km.

47

Because the three main volcanoes of Ross Island formed at different times, progressive deformation of the flexural surface is examined for Te = 4km. The flexural surface for the oldest, Mt Bird load, shows a quite circular pattern with maximum depth -

400m below Mt. Bird (Figure 68). The flexural surface becomes elongated in the NW-

SE direction in response to adding the Mt Terror load to the flexural deformation from

Mt. Bird (Figure 69). The maximum depth is -1000m below Mt. Terror. The flexural surface reaches a maximum depth of -1400m in response to adding the load of Mt

Erebus, summing the load from the entire Ross Island (Figure 70).

5.2.4 Best Fit Model

Because the depth of the pre-loading original surface is unknown, it is not meaningful to directly compare the depth of the model output with the depth of Rj from seismic observations. Instead, the dip angle of Rj is used to find the best fit model, with the assumption that the original surface before loading is a flat surface.

The Rj surface was mapped in 20 seismic lines. The velocity model (section

5.2.1) was used to convert Rj from two-way travel time to depth. Using this depth and over a defined horizontal distance, the average slope of the surface is calculated, and then the apparent dip angle of Rj is calculated using the arctangent of the average slope. The apparent dip angle calculated for the Rj surface ranges between 0.01-1.86 degrees, depending on the location and orientation of the seismic line (Table 5).

Depth to the flexural surfaces from the model output is extracted along the same track as the 20 seismic lines with Rj mapped. Based on the same procedure, the apparent

48 dip angle of the modeled surface is calculated for each seismic line for model outputs for

Te from 1 to 40 km.

For each Te, the 20 apparent dip angle from model output are compared with the observed apparent dip angle calculated in seismic lines. The root mean square (RMS) of the 20 misfits is calculated for each Te. The result (Figure 71) shows the minimum value of the RMS misfit is about 0.84 degree for Te of 4 km. The RMS of the misfit of apparent dips is almost the same when Te is 4 km and 5 km. So the best fit Te is 4-5km.

Using the model output for Te = 4km, the flexural surface resulting from the Mt.

Bird loading is compared with the Rj surface from seismic mapping (Figure 72) and the flexural surface of the model from only Mt. Erebus loading is compared with the Rk surface from seismic mapping (Figure 73). Since the direct comparison of depth is not that meaningful, the slope indicated by the density of contour lines is compared. The flexural surface modeled for Mt. Bird loading shows a circular pattern around Mt. Bird, similar to the observed concentric deepening of the Rj surface, with the exception of the deepening of the Rj surface toward the northwest, possibly due to overprint from Terror

Rift faulting. The contour density of the flexural surface indicates more gentle slopes than the mapped Rj surface west of Mt. Bird. The contour line density of the modeled flexural surface for Mt. Erebus loading fits reasonably well in shape and slope with the mapped

Rk surface around Mt. Erebus.

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Chapter 6 Discussion

6.1 Flexural Basin Evolution

6.1.1 Sequential Flexural Basins Related to Mt Bird and Mt Erebus Volcano

Loading

The Bird-Erebus flexural basin has a clear wedge-shaped fill that thickens toward the volcanoes of Ross Island, typical of the geometry of other basins formed by volcano loads in ocean basins. Seismic mapping in this study shows evidence of two discrete sub- basins within the flexural basin west of Ross Island. One is related to Mt. Bird loading with timing between the Rj and Rk surfaces, defined by Seismic Sequence 1 (SS1), and another is related to Mt. Erebus loading with timing later than the Rk surface, defined by

Seismic Sequence 2 (SS2). Various lines of evidence including the configuration of the seismic surfaces, unit isochore maps and the seismic facies and reflector geometry within units, are used below to document these two sub-basins.

6.1.1.1 Bird Sub-basin

Reflector Rj deepens towards Mt. Bird, indicating that the Rj surface is deformed due to Mt. Bird loading. The isochore map (Figure 52a) for SS1 between the Rj and Rk surface shows thickening in a partially concentric zone towards Mt. Bird, but this pattern is complicated by the linked zone of thickening extending northwest of Bird, discussed below as an overprint from Terror Rift. Seismic Units 1 and 2 both tilt and thicken

50 toward Mt Bird volcano, and reflectors proximal to Mt Bird have a divergent pattern, indicating that subsidence occurred throughout deposition of SS1 (Figure 74).

The zone of thicker deposits in SS1 around Mt Bird is interpreted to record deposition in a flexural moat basin isolated around the Mt Bird volcano. A flexural basin around Mt Bird has been predicted by prior modeling (Aitken et al., 2012), but has not previously been documented. The thinning to the south may record the edge of the Bird flexural basin. At least part of this pattern is related to erosion along the Rk unconformity, however, so the southward pinch out of SS1 is likely due to both flexure and glacial erosion.

In Unit 1 between the Rj and Rj1 reflectors, chaotic seismic facies interpreted as volcanogenic debris flow material extends around the border of Mt. Bird (Figure 47).

The volume and large spatial extent of the volcanic debris flow facies indicates that Unit

1 was deposited during the active growth phase of Mt. Bird. In Unit 2, patches of chaotic facies, interpreted as volcanic debris flow deposits, and possible distal volcanic-derived turbidite deposits, are present around Mt. Bird (Figure 47), recording ongoing mass wasting of the Mt Bird volcano, but with less activity than in Unit 1.

The main phase of Mt. Bird growth from onshore sample dating is from 3.8-4.6

Ma (Armstrong, 1978). The deflection of the lithosphere to form the flexural basin should form immediately following the main growth phase of the volcano. The Rj surface at the base of SS1 is below the volcanic apron of Mt. Bird, formed during the main phase eruption, indicating Rj is older than 3.8-4.6 Ma. This is older than the age for the Rj surface of about 3.2 Ma inferred within the AND-1B core south of Ross Island,

51 suggesting either the ‘jump correlation’ between that region and the basin west of Ross

Island is not correct, or that Rj is a time trangressive surface.

6.1.1.2 Erebus Sub-basin

The Rk seismic surface deepens towards Mt. Erebus (Figure 31a). This pattern indicates that the Rk surface deformed due to Mt. Erebus loading. The isochore map for

SS2 shows a zone of thickening toward Mt Erebus volcano. Although the circular pattern is less clear, due to seismic mapping limits imposed by the volcanic apron of Ross Island, the zone of thickening appears to be circumferential to Mt Erebus and to disappear northward in the basin adjacent to Mt Bird (Figure 52b). This zone is interpreted to mark flexural moat subsidence in response to loading by the Mt Erebus volcano. Again, although an Erebus flexural moat has been predicted (Aitken et al., 2012) and interpreted from a seismic grid of small spatial extent south of Mt Erebus (Horgan et al., 2005), documentation of a discrete basin is shown for the first time here.

In the basin around Mt Bird, seismic sequence 2 consists of a subhorizontal,

‘ponded’ unit that onlaps both basin margins and is thickest along the basin axis (Figure

74). This geometry indicates that there was passive accumulation of material within this sector of the basin, with no syn- or post-depositional tilting due to flexure. This geometry changes near the junction between Mt Bird and Mt Erebus. Closest to the junction, at Wolschlag Bay, seismic unit 2 thickens toward Mt Erebus, but strata within it remain horizontal (Figure 74). Further south, to the southwest of Mt Erebus, seismic sequence 2 both thickens and tilts toward Mt Erebus, and has divergent internal reflector patterns, indicating that flexure was occurring in this zone during deposition of the

52 sequence (Figure 74). The seismic surface maps, the isochore maps and the internal reflector geometries for the units that make up SS2 show that flexure stopped in the Bird sub-basin by Unit 3 time, then stopped to the northwest of Erebus in Unit 4 time, when flexure was still ongoing only in the area to the southwest of Erebus.

In seismic Unit 3 and Unit 4, chaotic seismic facies interpreted to be volcanogenic debris flow deposits occur along the west and northwest flanks of Mt. Erebus (Figure 49,

Figure 50). Seismic Unit 5 consists of an up to 100 ms thick volcanic debris flow with large spatial extent southwest of Mt. Erebus (Figure 51). The abundance of volcanic debris flow deposits around Mt Erebus indicates these units were deposited at the same time or after Mt. Erebus formed, filling the sub-basin associated with Mt. Erebus loading.

The main phase of Mt. Erebus from onshore sample dating is constrained to have started about 1.34 Ma, which is interpreted to mark the transition from the subaqueous to terrestrial shield-building phase of the volcano, and the proto-Erebus shield volcano was formed by ~0.75 Ma (Esser et al., 2004). The sub-basin associated with Mt Erebus loading should form immediately following the main phase eruption. The Rk surface formed earlier than the major volcanic debris flows around the west side of Mt. Erebus, so the Rk surface there is likely older than 1.34 Ma. The Rk reflector in the AND-1B core south of Ross Island is dated at 2 Ma (Wilson et al., 2012). This ‘jump correlation’ from at 2.0 Ma Rk across to a >1.34 Ma Rk in the basin west of Erebus is not inconsistent, although it is a significant age difference given the rapid volcano growth.

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6.1.2 Relation to Terror Rift

The northwest-trending zone marked by deepening of seismic reflectors Rj, Rj1 and Rk, and by thickening in isochore maps of Units U1, U2 and U3, is spatially associated with normal faults that define the Terror Rift (Figures 52a, Figure 52b). This zone is interpreted to represent deposition in accommodation space created by rift faulting. The link between the fault-related zone and the concentric accommodation space surrounding Mt Bird indicates that flexure due to volcano loading was happening at the same time as Terror Rift stretching. Considering the isochore maps for the individual seismic units, it appears that Terror Rift activity may have ended by the time that U4 was deposited, as it has relatively constant thickness across the region (Figure 32b). U4 is quite thin in the northern basin region, however, and may have been partly removed by glacial erosion, so younger rift activity can’t be ruled out.

6.1.3 Infilling of Flexural Basin

6.1.3.1 Sedimentary Record in Bird-Erebus Flexural Basin west of Ross Island

Modern day seafloor bathymetry shows a deep flexural moat around Ross Island with seafloor depths typically exceeding 800 mbsl (Figure 2), except the area southwest of Mt. Erebus, where it shoals to ~500 mbsl. The majority of the flexural basin still has large accommodation space, which means the flexural basin is underfilled. Most of the flexural basin was likely underfilled throughout its development, with a total of only

~325 m of sedimentary fill accumulating over most of the basin (Figure 53). Thicker fill only occurs in the basin to the southwest of Mt Erebus, which filled with more volcanic

54 material derived from collapses of the flank or summit calderas, which occurred in the later stages of Erebus volcano growth (Esser et al., 2004; Harpel et al., 2004).

In AND-1B core, the dominant lithologies in Pliocene strata above the Rj surface are diamictite, deposited in subglacial to ice-proximal glacial marine environments, and diatomite deposited in an open marine environment (Naish et al. 2007). In Pleistocene strata sampled by the core above the Rk surface, open-marine deposits are mudstones and diatomites, subglacial diamictites are common, and volcaniclastic material is a significant component (McKay et al., 2012). The Plio-Pleistocene deposits are interpreted to record repeated ice sheet advance/retreat cycles with input of volcanic material from active centers on Ross Island and to the south. These lithologies also form a major component in the flexural basin to the west of Ross Island, as recorded in particular by the widespread occurrence of the seismic F1 laminated facies (Figure 47).

Seismic facies distribution maps (Figure 47-52) show that sediment types have high spatial variation, and change though time. The components of volcanogenic gravity flow material, including volcanic debris flows and volcaniclastic turbidites, increase spatially closer to the volcanoes. There are more chaotic units, interpreted as debris flows, input into the basin during the active construction stage than the post-growth mass wasting stage of the volcanoes, first around Mt Bird (Figure 47, Figure 48) and later around Mt Erebus (Figure 49, Figure 50, and Figure 51).

The sedimentary fill of the flexural basin has clearly been affected by glacial erosion. Numerous ‘glacial surfaces of erosion’, interpreted to mark advance of grounded ice during ice sheet advances, have been interpreted in the AND-1B core

55 sequence (Naish et al., 2009; McKay et al., 2009) that is equivalent to the strata imaged seismically in this study. Clearly the Rk seismic surface has significant erosional relief across the whole basin and is a regionally important glacial erosion surface where strata were removed. Although the amount of erosion can’t be easily quantified, it is important to note some of the sedimentary record of the Bird-Erebus flexural basin has certainly been lost during ice sheet advance/retreat cycles.

6.1.3.2 Comparison with Sedimentary Record of other Flexural Basins formed by

Volcano Loading

Seismic stratigraphy work from flexural moat basins around the Hawaiian Islands

(Rees et al., 1993), Marquesas Islands (Wolfe et al., 1994), Canary Islands (Collier and

Watts, 2001), and Cape Verde Islands (Ali et al., 2003) have provided information on how flexural basins form and how they are infilled with time. All of these flexural basins formed in relatively deep water, on oceanic crust, with the Hawaiian and Marquesas sited in the deep ocean basin and Cape Verde and Canary Islands off the continental margin of

Africa. None of the basin sequences have been drilled to obtain lithology information directly, so seismic facies and velocities from seismic reflection profiles have been used to infer sediment sources. In all the basins, high seismic velocities and abundant units with internally chaotic seismic facies are correlated with a higher proportion of volcanic material derived from the volcanic islands by mass wasting processes.

The Marquesas and Hawaii flexural basins are mostly filled with material derived from gravitational mass wasting of volcanoes. These two basins have fill dominated by chaotic units, with seismic velocities in the central moat up to > 4000 m/s at Marquesas,

56 and 3700 -4700 m/s at Hawaii. Marquesas is overfilled, with an extensive volcanic apron that completely fills and covers the flexural moat, interpreted to indicate repeated collapse and build up of the volcano summits, producing large volumes of debris that accumulated in the central basin. The flexural moat around the Hawaiian chain changes from nearly filled next to the older volcanic islands to underfilled near the youngest, a pattern attributed to the time span available for mass wasting of the volcanoes to supply debris to the basin (Rees et al., 1993).

The flexural basins around the Cape Verde Islands and Canary Islands are also filled largely with volcanic material, but components of continental margin derived materials also reached these basins. The seismic facies in the flexural moats of Cape

Verde Islands and Canary Islands looks less chaotic than Hawaii and Marquesas. Seismic velocities of flexural moat fill is 2500 -3200 m/s at Cape Verde Islands and 2000-3000 m/s at Canary Islands. Their velocities are lower than in the material infilling the

Marquesas and Hawaii basins.

Compared with other flexural sedimentary basins, the Bird-Erebus basin has a few different characteristics. Although the basin clearly has a substantial component of volcanogenic fill related to mass wasting of the volcanic islands, it does not appear to be the dominant fill as interpreted elsewhere. Instead, the Bird-Erebus basin has a significant component of glaciomarine deposits and, based on correlation with the AND-

1B core, deposits that record repeated advance and retreat of a marine-based ice sheet across the region. The low seismic velocities for the basin sediments are consistent with a relatively small component of volcanic infill compared to the other flexural moat

57 basins. The glacial cycles also resulted in erosion of parts of the sedimentary record.

Erosion of basin fill is not a significant factor in the other flexural basins. Rees et al.

(1993) emphasized that for the Hawaiian flexural moat basin, and probably others, there was unlikely to be a continuous stratigraphic record through the life of the evolving flexural basin because volcanic mass wasting processes that filled the basin were episodic. In the Bird-Erebus basin, the record is likely to be even more discontinuous due to episodic volcanogenic input, as well as episodic deposition and erosion during glacial cycles.

6.1.4 Basin Fill Geometry and Progressive Volcano Loading

Seismic stratigraphic studies of the Hawaiian flexural basin documented a distinct stratigraphic pattern of the infill, with a) early sequences having wedge shapes, internal divergent reflectors, and showing onlap away from the island loads and toward the flexural bulge, b) younger sequences showing tilting but offlap back toward the island loads, and c) the uppermost strata ‘ponded’ in the deepest sector of the moat (ten Brink and Watts, 1985; Rees et al., 1993). The same overall pattern occurs in the Marquesas

(Wolfe et al., 1994), Canary (Collier and Watts, 2001), and Cape Verde (Ali et al., 2003) flexural moat basins, all linear island chains where volcanoes formed at different times.

In both the Canary and Cape Verde examples, concentric thickening of the basin deposits occurrs around the volcanoes that caused the loading.

Watts and ten Brink (1989) modeled a flexural basin with progressive loading of a volcanic island chain on an elastic plate, to simulate the formation of the Hawaiian flexural moat. Their model included sediments that continuously filled the basin to a 58 reference level. This elastic plate progressive loading model reproduces the cross-moat pattern of onlap followed by offlap, and also predicts strata tilting toward the new load in profiles of the basin along the direction of progressive volcano loading, onlapping in the lower section and offlapping in the upper section (Figure 75). This model prediction was verified by seismic mapping by Rees et al. (1993).

In the Bird-Erebus flexural basin, concentric thickening toward the older Bird volcano and younger Erebus volcano occurs in successive seismic sequences. The typical wedge geometry in cross-moat profiles is present. However, the entire stratal sequence is characterized by onlap geometries in cross-moat profiles, consisting of onlap away from the islands in the lower sequence and onlap over both basin margins in the upper sections. No offlap pattern is observed in cross-moat or along moat profiles. Either offlapping units did not form, or were completely removed by erosion. Seismic lines that are near-parallel to the Bird-Erebus flexural basin show that there is a regional northward tilt to the strata, not the tilt south toward the younger Mt Erebus volcano that the progressive load model predicts.

6.2 Flexural Modeling and Lithosphere Strength

6.2.1 Comparison with Previous Ross Island Flexural Modeling

Aitken et al. (2012) presented results of 3D flexural modeling of progressive loading by all the major volcanoes of the Erebus Volcanic Province, using constraints from seismic and gravity profiles from south of Ross Island. Their results show the best fit Te is between 2-5 km, selecting 3 km as the elastic thickness that best fits all

59 constraints. This Te is similar to the best fit Te of 4-5 km from this study, which is constrained by a more extensive set of seismic lines directly adjacent to the western margin of the Ross Island volcanoes.

When Te = 3km, the lithosphere forms two main regional basins in response to progressive loading by volcanoes of the Erebus Volcanic Province, one around southern volcanoes and one centered at Ross Island (Aitken et al., 2012) (Figure 76). The deformation pattern of progressive loading from Mt. Bird, Mt. Terror, and Mt. Erebus produced in the modeling by Aitken et al. (2012) is similar to the flexural pattern modeled in this work. The modeled basin widths are more narrow and the magnitude of predicted subsidence is greater for the Aitken et al. (2012) model, due to the lower preferred Te value and the assumption that sediment fills all accommodation space between emplacement of each successive volcano, adding to the total load. The orientation of the flexural basins formed after each Ross Island volcano load has a northeast-southwest elongation in the Aitken et al. (2012) results, whereas model results here predict a circular basin from the Mt Bird volcano load, and a basin with an overall northwest-southeast trend after Mt Terror and Mt Erebus volcano loads are added. This difference may be due to how the loading grid is cropped from the DEM in the two studies; the Aitken et al. (2012) modeling used rectangular loading grids whereas this study used loading grids following the subsurface apron of each volcano.

Figure 77 from Aitken et al. (2012) shows the predicted stratigraphy of the final flexural basin below Ross Island along a profile going through the summits of Mt. Erebus and Mt. Terror, produced with the model assumption that accommodation space created

60 by each load filled to the -500 m reference level before the next load started. This prediction can be compared to the stratigraphy from seismic line NBP0401-157 (Figure

74) from this study. The stratigraphic pattern in terms of dip and relative stratigraphic thicknesses is similar between the modeled accommodation space and the mapped seismic stratigraphy. There is only a thin and spatially limited sequence from the Mt Bird load in this part of the basin, and a wedge-shaped, eastward-thickening sequence related to the Mt Erebus load. The seismic mapping completed in this study could not differentiate a sequence related to Mt. Terror load. The overall thickness of the modeled stratigraphy of Aitken et al. (2012) is much thicker than the actual sedimentary fill mapped in this work. The present study shows that the flexural basin was underfilled, and that sedimentary material was removed by glacial erosion, resulting in the stratal thicknesses lower than predicted by the model.

Modeling by Aitken et al. (2012) and in this study both indicate that a very low

Te best fits available constraints. This weak lithosphere in Ross Island region can be explained by the rift tectonic setting. The development of Terror Rift north of, and probably beneath, Ross Island causes faulting (Cooper et al., 1987; Henrys et al., 2007), thin crust (Bannister et al., 2003; Lawrence et al., 2006), and high heat flow (Blackman et al., 1987; Morin et al., 2010; Schroder et al., 2011) in the Ross Island region. The seismically-mapped stratal thicknesses defined by this study show that the Terror Rift was active at the same time that volcano loading was deforming the lithosphere. Stern et al. (1991) applied a 2D broken plate model to study the flexural loading by the Ross

Island volcanoes. The circular pattern of the flexural basin, shown by both modern day

61 basin bathymetry and sediment thickness patterns from seismic mapping, is most compatible with a continuous elastic plate model, rather than a broken plate model.

6.2.2 Flexural Modeling Results Compared with Other Basins Formed by Volcanic

Island Loading

The flexural response of many intraplate volcanic ocean islands has been examined, such as the Hawaiian Islands (Watts and ten Brink, 1989; Rees et al., 1993),

Canary Islands (Collier and Watts, 2001), and Cape Verde Islands (Ali and Watts, 2001).

Each of these basins formed by progressive emplacement of volcanoes along linear zones, resulting in composite flexural basins. In general, the size of the ocean island volcano loads in these regions is larger than Ross Island, and the surrounding flexural moats and flexural bulges are deeper and wider. For example, the volume of volcanic loading from the Hawaiian Islands is about 2 × 1014 푚3, which is about 44 times that of

Ross Island. The flexural moat, with radius of about 140 km, forms a semicircular trough around the southeastern end of the Hawaii Islands, has ~1,000 m in relative relief with the flexural bulge, which is located about 250 km away from the islands. Modeling results yield lithospheric strengths, represented by Te, that are much stronger than found for the Ross Island area. Modeled Te is 25-30 km for the Hawaiian Islands, 35 km for the Canary Islands, and 29 km for the Cape Verde Islands. All of these flexural basins have total sedimentary infills over 2,000 m thick, compared with the ~325 m fill of the

Bird-Erebus flexural baisn fill. Even with the stronger lithosphere, however, the Canary and Cape Verde flexural basins show concentric thickening around volcanoes related to progressive loading, as found in this study for the Mt Bird and Mt Erebus volcano loads. 62

Chapter 7: Conclusions

Seismic reflection data located west of Ross Island are used here to map the seismic stratigraphy of the Bird-Erebus flexural basin around Ross Island. Five seismic surfaces are mapped within the basin sequence, bounding 5 seismic units. The Rj seismic surface lies below the volcanic apron of the Mount Bird volcano, defining the base of the flexural basin infill. The Rk seismic surface formed prior to flexure after Mount Erebus volcano growth, and shows substantial incision into the underlying unit, documenting erosion due to ice sheet advance over the region. Seven seismic facies are identified based on seismic reflection characteristics, and are interpreted as volcanic sediment gravity flows from Ross Island, subglacial deposits, glaciomarine deposits, and open marine deposits infilling the basin.

The five seismic units are categorized into 2 seismic sequences. Seismic Sequence

1 consists of the 2 units between the Rj and Rk surfaces, which dip toward Mt Bird and show increasing thickness in a roughly concentric zone around Mt Bird. Seismic

Sequence 2 is made up of the 3 seismic units between the Rk surface and the seafloor.

These units dip and thicken toward Mt Erebus. Each seismic sequence shows a general pattern that includes seismic surface dips toward the related volcano load, thickness increasing toward the load, internal seismic reflector divergence toward the load, and chaotic seismic units interpreted as volcanic mass flows localized around the volcano load. The sequence patterns are interpreted to record two discrete sub-basins formed by

63

Mt. Bird loading at 4.6-3.8Ma and Mt Erebus loading after ~1.31 Ma. Both sequences also show a zone of increased depth and isochore thickness with a linear western boundary that extends northwest from Mt Bird. This zone is spatially coincident with

Terror Rift faults and is interpreted to represent deposition in rift accommodation space at the same time as flexural deformation was taking place.

The strata in the Bird-Erebus flexural basin have an overall wedge-shaped fill that thickens toward the volcanoes of Ross Island, typical of the geometry of other flexural moats formed by volcano loads in ocean basins. Previous seismic stratigraphy and modeling of progressive loading in ocean island chains shows that strata onlap in the lower section of basin strata and offlap in the upper section in both cross-moat and along moat profiles (ten Brink and Watts, 1985; Collier and Watts, 2001; Ali et al., 2003). Ross

Island flexural basin strata show onlap geometries in cross-moat profiles. However, no offlap pattern is observed in either cross-moat or along moat profiles.

The flexural basin is underfilled, with a flexural basin fill of only ~325 m around western and southern Ross Island, compared with over 2000 m of fill in other oceanic flexural moat basins. Sediment gravity flow deposits related to mass wasting of Ross

Island volcanoes forms aprons in the basin fill along the volcano slopes. However, unlike other basins formed by volcano loads, volcanogenic fill related to mass wasting of the volcanic islands is not the dominant fill in the Ross Island flexural basin. Instead, seismic facies mapping and AND-1B core show a significant component of glaciomarine, subglacial and open marine deposits in the flexural basin.

64

3D thin elastic plate modeling used a velocity model based on data from the

AND-1B core to convert two-way travel time to depth for seismic surface Rj. Using the dip angle of the Rj surface as the constraint, the best fit lithospheric strength is represented by an effective elastic thickness (Te) of 4-5 km. This result is consistent with a previous 3D modeling study of the Ross Island region by Aitken et al. (2012), which showed a best fit Te of 2-5 km. The lithosphere in the study area is significantly weaker than the oceanic lithosphere where volcanic loading has been modelled elsewhere (Watts and ten Brink, 1989, Collier and Watts, 2001, Ali et al., 2003). This weak lithosphere in the Ross Island region can be explained by the rift tectonic setting. The development of

Terror Rift north of, and probably beneath, Ross Island is associated with faulting

(Cooper et al., 1987; Henrys et al., 2007), thin crust (Bannister et al., 2003; Lawrence et al., 2006), and high heat flow (Blackman et al., 1987; Morin et al., 2010; Schroder et al.,

2011) in the Ross Island region.

65

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Vail, P. R. (1987). Seismic stratigraphy interpretation using sequence stratigraphy: Part 1: Seismic stratigraphy interpretation procedure. Veeken, P. P. (2013). Seismic Stratigraphy and Depositional Facies Models. Academic Press. Walcott, R. I. (1970). Flexural rigidity, thickness, and viscosity of the lithosphere. Journal of Geophysical Research, 75(20), 3941-3954. Watts, A. B. (2001). Isostasy and Flexure of the Lithosphere. Cambridge University Press. Watts, A. B., & Ten Brink, U. S. (1989). Crustal structure, flexure, and subsidence history of the Hawaiian Islands. Journal of Geophysical Research: Solid Earth (1978– 2012), 94(B8), 10473-10500. Watts, A. B., & Burov, E. B. (2003). Lithospheric strength and its relationship to the elastic and seismogenic layer thickness. Earth and Planetary Science Letters, 213(1), 113-131. Watts, A. B., Ten Brink, U. S., Buhl, P., & Brocher, T. M. (1985). A multichannel seismic study of lithospheric flexure across the Hawaiian–Emperor seamount chain. Nature, 315(6015), 105-111. Watson, T., Nyblade, A., Wiens, D. A., Anandakrishnan, S., Benoit, M., Shore, P. J., ... & VanDecar, J. (2006). P and S velocity structure of the upper mantle beneath the Transantarctic Mountains, East Antarctic craton, and Ross Sea from travel time tomography. Geochemistry, Geophysics, Geosystems, 7(7). Wessel, P., Smith, W. H., Scharroo, R., Luis, J., & Wobbe, F. (2013). Generic mapping tools: improved version released. Eos, Transactions American Geophysical Union, 94(45), 409-410. Widess, M. B. (1973). How thin is a thin bed?. Geophysics, 38(6), 1176-1180. Wilson, T. J. (2004). Cruise Report NBP0401, 19 January to 18 February 2004, McMurdo Station to McMurdo Station, Ross Sea, Antarctica. Institute of Geological & Nuclear Sciences. Wilson, G. S., Levy, R. H., Naish, T. R., Powell, R. D., Florindo, F., Ohneiser, C., ... & Pollard, D. (2012). Neogene tectonic and climatic evolution of the Western Ross Sea, Antarctica—Chronology of events from the AND-1B drill hole. Global and Planetary Change, 96, 189-203. Whittaker, J. (2005). Cenozoic structural and stratigraphic history of McMurdo Sound, Antarctica. Master's Thesis, Victoria University of Wellington, Wellington, New Zealand. 71 pp.

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Appendix A: Tables

Inferred Fielding Horgan Bartek Brancolini Cooper Nature of Boundaries Whittaker Aged This study Tectonic Stage et al. et al. et al. et al. et al. and Facies of Units Thesis (This study) (2008) (2005) (1996) (1995) (1987) Seafloor F5 U5 Broadly undulating Rk2 boundary.Truncate U4 F1, F3 and F5 U4 SUB Mt. Erebus forming Smooth, mostly Rk1 (Seismic Sequence conformable boundary. 2 in this study) F2, F3, and F5 U3

Rough erosional boundary A-B Surface > 1.34 Ma with sharp incisions. Rk H2 Rk A0 Truncate U1 and U2. F1, F2, F3, F4, F5, and F6 U2 Broadly undulating V1 boundary. Conformable Rj1 Mt. Bird forming SUC towards north, truncating (Seismic Sequence U1 towards south 1 in this study) F1, F2, F3, F5, and F6 U1 Smooth, conformable Surface > 3.8-4.6 Ma Rj H3 Rj C-D? RSU1? boundary. A1 SUD H4 SUE Surface 4.5Ma H5 Ri E-F RSU 2 A2 RSS 5-8 VLB SUF G-H RSU 4 Rift Stage 5 RSS 4 Surface 7.5Ma H6 Rh RSU 4a B SUG I, J, K RSS 3 V2 Surface 11Ma H7 Rg L-M RSU 5 C SUH VLB H8 Rift Stage 4 SUI N-P RSS 2 V3 H9 SUJ 17 Ma H10 Rf P/Q ~RSU 6 Table 1: Seismic stratigraphic framework for this study.

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Example Images Facies Occurrence Vertical Lateral Amplitude Internal External Facies Interpretation Name Frequency Continuity Reflector Shape Geometry F1 High High High High Parallel Sheet Alternating diamictite and biogenic or hemipelagic; volcaniclastic layers

Subglacial, glaciomarine & open marine F2 High Low Moderate to Moderate to Sub- Sheet to high low low parallel Diamictites, or pelagic/hemipelagic/sand with ice-rafted debris

Glacial marine, or subglacial deposit

F3 High High Moderate Moderate to Lenticular Channel high form, lens, Sediment gravity flows, possible bottom current or mound reworking;

Glaciomarine

F4 Low High High Moderate Oblique Truncated parallel clinoform? Alternating sediment gravity flows & Or tilted? hemipelagic?

Deep water current or slope fan ? F5 High Low Low Moderate to Chaotic Sheet, or high with point mound Volcanogenic sedimentary mass flows; source diamictites diffraction Volcanic apron; subglacial to ice- proximal glaciomarine F6 Medium Low Low Low Reflection Sheet free Homogeneous diatomite, mud or sand

Glaciomarine deposits

F7 Medium Low Low Low Chaotic Steep mound Volcanic body

Table 2: List of 7 seismic facies with their reflection characteristics and possible interpretations.

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Highest Volume calculated Diameter Volume in this Loads elevation by Esser et al. (2004) (km) study (풎ퟑ) above -1200m (풎ퟑ)

Mt. Bird 2984.99 15.285 9.256 × 1011 4.701 × 1011

Mt. Terror 4479.28 21.756 3.238 × 1012 1.796 × 1012

Mt. Erebus 4981.71 21.366 2.960 × 1012 2.253× 1012

Ross Island total 7.124 × 1012 4.519 × 1012

Table 3: Volume of Ross Island volcanos.

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Parameter Meaning Value Unit

3 휌푙 Density of loading material 2650 kg/m

3 휌푤 Density of water 1030 kg/m

3 휌푚 Density of mantle 3260 kg/m

3 휌푖 Density of infill material 1030 kg/m

g Gravitational acceleration 9.806199203 m/s2

E Young’s Modulus 1011 Pa

ν Poisson’s Ratio 0.25 N/A

Te Effective Elastic Thickness 1, 2, 3,….., 39, 40 km

Table 4: Parameters used in flexural model.

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Seismic lines Rj dip angle (Degree) NBP0401-100 1.862648 NBP0401-126 1.292978 NBP0401-128 1.462411 NBP0401-130m 0.530406 NBP0401-151d 1.168977 NBP0401-151e 0.017701 NBP0401-151g 0.488846 NBP0401-152 1.616485 NBP0401-157 0.875669 NBP0401-158 1.527775 NBP0401-159 1.199718 IT90a-70S 0.834975 IT90a-69 1.424937 IT90a-70N 1.262831 IT90a-70NN 0.9354 IT90a-73 1.760597 IT90a-74 0.460412 hpp2 1.79806 mis1 0.170923 mis2 0.754076

Table 5: Dip angle of seismic reflector Rj in 20 seismic profiles.

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Appendix B: Figures

Figure 1. Upper: Structure of West Antarctic Rift System in the Ross Sea; green denotes Terror Rift within the Victoria Land Basin; Lower: Faulting associated with the southern Terror Rift in the McMurdo Sound area. From Henrys et al. (2007). 78 Figure 2. Bedrock topography of McMurdo Sound. Onshore elevations denoted by yellow – white color ramp, bathymetric elevations denoted by blue color ramp. Black lines show seismic data coverage. Yellow dot shows location of AND-1B core.

79 seamoun t 1.94 Ma

Figure 3 Volcano ages in the Erebus Volcanic Province modified from DiRoberto et al. (2012); age and approximate extent of volcanic seamount east of Beaufort Island from Rilling et al. (2009).

80 Figure 4. Crustal thickness from the McMurdo Sound and surrounding region. Circle values are crustal thickness (km) from teleseismic data (Bannister et al., 2003); box values are crustal thickness (km) from multichannel seismic reflection data (McGinnis et al., 1985; Cooper et al., 1997).

81 Figure 5. Heat flow values for the McMurdo Sound region in mW m2 (Blackman et al., 1987; Morin et al., 2010; Schroder et al., 2011).

82 Figure 6. Airy-Heiskanen at continental scale (from Watts, 2001).

83 Figure 7. Vening-Meinesz's isostasy model. From Wienecke (2006).

84 Figure 8. Schematic illustrating multibeam sonar survey. Survey vessel is located at sea level, and swath width denotes sea floor coverage of sonar measurements (L-3 Communications, 2000).

85 Figure 9. Schematic illustrating multichannel marine seismic reflection survey (http://openlearn.open.ac.uk/mod/resource/view.php?id=172129)

86 Figure 10. Workflow for this study. Orange boxes represent steps associated with seismic data analysis and interpretation. Green boxes represent steps associated with modelling.

87 Figure 11. Schematic of seismic sequence boundary illustrating the geometric relationship between reflectors (red arrows) and upper and lower boundaries (green lines) (Veeken and Moerkerken, 2013 ).

88 Figure 12. General reflection configurations (Schlaf et al., 2005)

89 Figure 13. Reflection configurations of clinoforms (Schlaf et al., 2005)

90 Figure 14. Seafloor multiples from seismic line NBP0401-118, located within the McMurdo Sound region.

91 Figure 15. Seismic reflector pull up due to igneous body from seismic line NBP0401-158.

92 Figure 16. Four geographic regions defined in McMurdo Sound, west of Ross Island. White lines with numbers are contour lines. Summit height of volcanoes are labeled.

93 F1, example from NBP0401_128:

F2, example from IT90a_74:

F3, example from NBP0401_151g:

F4, example from NBP0401_151d:

Figure 17. Examples of seismic facies: F1 to F4. Horizontal distance is about 2km (Blue line) Vertical scale of 100 ms is set at F1 (Blue line). See Table 2 for description and interpretation. 94 F5, example from NBP0401_119:

F6, example from IT90a_69:

F7, example from NBP0401_151e:

Figure 18. Examples of seismic facies: F5 to F7. Horizontal distance is about 2km (Blue line) Vertical scale of 100 ms is set at F1 (Blue line). See Table 2 for description and interpretation.

95 Figure 19. Spatial distribution of Seismic Facies F1 (in green). Basemap shows the modern shelf-slope-basin bathymetry (contours) and channel systems (bold arrows) mapped by Stutz (2012).

96 Figure 20. Lateral seismic facies transition from F2 to F6 in seismic line IT90a-70.

97 Figure 21. Spatial distribution of Seismic Facies F2 (in blue). Basemap shows the modern shelf-slope-basin bathymetry (contours) and channel systems (bold arrows) mapped by Stutz (2012).

98 Figure 22. Spatial distribution of Seismic Facies F3 (in yellow). Basemap shows the modern shelf-slope-basin bathymetry (contours) and channel systems (bold arrows) mapped by Stutz (2012).

99 Figure 23: Spatial distribution of Seismic Facies F4 (in gray with pink outline). Basemap shows the modern shelf-slope-basin bathymetry (contours) and channel systems (bold arrows) mapped by Stutz (2012).

100 Figure 24. Spatial distribution of Seismic Facies F5 (in purple). Basemap shows the modern shelf-slope-basin bathymetry (contours) and channel systems (bold arrows) mapped by Stutz (2012).

101 Figure 25. Spatial distribution of Seismic Facies F6 (in pink). Basemap shows the modern shelf-slope-basin bathymetry (contours) and channel systems (bold arrows) mapped by Stutz (2012).

102 Figure 26. Spatial distribution of Seismic Facies F7 (in red). Basemap shows the modern shelf-slope-basin bathymetry (contours) and channel systems (bold arrows) mapped by Stutz (2012).

103 Figure 27. Alternation of diamictite (green lithology) and diatomite (yellow lithology) and associated seismic reflection, velocity, and density properties at AND-1B core (Hansaraj, 2008).

104 Figure 28. Interglacial sedimentation model in McMurdo Sound, Antarctica (Bartek and Anderson, 1991).

105 a b

Figure 29. a) Rj seismic surface depth in two-way travel time. b) Isochore map of U1 seismic unit thickness in two-way travel time.

106 a b

Figure 30. a) Rj1 seismic surface depth in two-way travel time. b) Isochore map of U2 seismic unit thickness in two-way travel time.

107 a b

Figure 31. a) Rk seismic surface depth in two-way travel time. b) Isochore map of U3 seismic unit thickness in two-way travel time.

108 a b

Figure 32. a) Rk1seismic surface depth in two-way travel time. b) Isochore map of U4 seismic unit thickness in two-way travel time.

109 a b

Figure 33. a) Rk2 seismic surface depth in two-way travel time. b) Isochore map of U5 seismic unit thickness in two-way travel time.

110 Figure 34. Location of seismic lines described in Figures 35 – 46 and discussed in text section 5.1.2: Description of Seismic Surfaces and Units.

111 Figure 35. Seismic line NBP0401-100. Colored lines denote interpreted seismic surfaces. Bold x markers indicate intersection of two seismic profiles. Convex reflector segments (pink) represent volcanic bodies.

112 Figure 36. Seismic line NBP0401-151d. Colored lines denote interpreted seismic surfaces. Bold x markers indicate intersection of two seismic profiles. Convex reflector segment (pink) represents volcanic body. Seismic line intersects with Mt. Bird apron to the south (right of figure).

113 Figure 37. Seismic line IT90a-70. Colored lines denote interpreted seismic surfaces. Bold x markers indicate intersection of two seismic profiles. Convex reflector segment (pink) represents volcanic body.

114 Figure 38. Seismic line IT90a-74. Colored lines denote interpreted seismic surfaces. Bold x markers indicate intersection of two seismic profiles.

115 Figure 39. Seismic line IT90a-73. Colored lines denote interpreted seismic surfaces. Bold x markers indicate intersection of two seismic profiles. Convex reflector segments (pink) represent volcanic bodies.

116 Figure 40. Seismic line NBP0401-128. Colored lines denote interpreted seismic surfaces. Bold x markers indicate intersection of two seismic profiles. Seismic line intersects with Mt. Bird apron to the east (right of figure).

117 Figure 41. Seismic line NBP0401-126. Colored lines denote interpreted seismic surfaces. Bold x markers indicate intersection of two seismic profiles. Convex reflector segment (red) represents volcanic body.

118 Figure 42. Seismic line NBP0401-159. Colored lines denote interpreted seismic surfaces. Bold x markers indicate intersection of two seismic profiles.

119 Figure 43. Seismic line NBP0401-157. Colored lines denote interpreted seismic surfaces. Bold x markers indicate intersection of two seismic profiles. Convex reflector segment (pink) represents volcanic body. Seismic line intersects with Mt. Erebus apron to the southeast (right of figure).

120 Figure 44. Seismic line NBP0401-158. Colored lines denote interpreted seismic surfaces. Bold x markers indicate intersection of two seismic profiles. Convex reflector segment (red) represents volcanic body.

121 Figure 45. Seismic line NBP0401-130m. Colored lines denote interpreted seismic surfaces. Bold x markers indicate intersection of two seismic profiles.

122 Figure 46. Seismic line NBP0401-125. Colored lines denote interpreted seismic surfaces. Bold x markers indicate intersection of two seismic profiles. Convex reflector segment (pink) represents volcanic body.

123 Figure 47. Seismic facies distribution within seismic unit U1 (extent of unit shown as pink polygon). Basemap shows the modern shelf-slope-basin bathymetry (contours) and channel systems (bold arrows) mapped by Stutz (2012). Green arrow indicates north.

124 Figure 48. Seismic facies distribution within seismic unit U2 (extent of unit shown as pink polygon). Basemap shows the modern shelf-slope-basin bathymetry (contours) and channel systems (bold arrows) mapped by Stutz (2012). Green arrow indicates north.

125 Figure 49. Seismic facies distribution within seismic unit U3 (extent of unit shown as pink polygon). Basemap shows the modern shelf-slope-basin bathymetry (contours) and channel systems (bold arrows) mapped by Stutz (2012). Green arrow indicates north.

126 Figure 50. Seismic facies distribution within seismic unit U4 (extent of unit shown as pink polygon). Basemap shows the modern shelf-slope-basin bathymetry (contours) and channel systems (bold arrows) mapped by Stutz (2012). Green arrow indicates north.

127 Figure 51. Seismic facies distribution within seismic unit U5 (extent of unit shown as pink polygon). Basemap shows the modern shelf-slope-basin bathymetry (contours) and channel systems (bold arrows) mapped by Stutz (2012). Green arrow indicates north.

128 a b

Figure 52. a) Isochore of Seismic Sequence 1. b) Isochore of Seismic Sequence 2. Seismic unit thicknesses (red – purple color scale) are in two-way travel time. Terror Rift normal faults north of Mt. Bird designated by bold gray fault lines (from Henrys et al., 2007).

129 Figure 53. Isochore of Flexural Basin Sequence from Rj to seafloor. Seismic unit thickness (red – purple color scale ) is two-way travel time.

130 Figure 54. Location of seismic lines described in Figures 55 – 61 and discussed in text section 5.1.4: Age Constraints for Seismic Surfaces and Units.

131 Figure 55. Seismic line NBP0401-151g. Colored lines denote interpreted seismic surfaces. Orange arrows indicate pattern of sediment on lapping onto Rj1 surface (top of Mt. Bird apron). Reflector Rj is below the Mt. Bird apron.

132 Figure 56. Seismic line IT90a-70. Colored lines denote interpreted seismic surfaces. Bold x markers indicate intersection of two seismic profiles. Convex reflector segments (red) represent volcanic body.

133 Figure 57. Geomorphology and age data show Mt. Erebus collapse calderas. Black arrow shows direction of volcanogenic mass flow.

134 Figure 58. FR1 in seismic line NBP0401-126. Colored lines denote interpreted seismic surfaces. The body enclosed inside of red line represents fissure ridge. Rj, Rj1, Rk and Rk1 extend below fissure ridge.

135 Figure 59. FR2 in seismic line NBP0401-157. Colored lines denote interpreted seismic surfaces. The body enclosed inside of red line represents fissure ridge. Rj and Rk extend below fissure ridge, Rk1 terminates at fissure ridge.

136 Figure 60. FR3 in seismic line NBP0401-158. Colored lines denote interpreted seismic surfaces. The body enclosed inside of red line represents fissure ridge. Rj and Rk extend below fissure ridge, Rk1 terminates at fissure ridge.

137 Figure 61. FR4 in seismic line NBP0401-123. The body enclosed inside of red line represents fissure ridge. Rk and Rk1 extend below fissure ridge.

138 Figure 62. Age constraints of seismic surfaces around AND-1B drillsite. Left section is seismic surfaces mapped by Hansaraj (2008). Right section is chronology data from AND-1B core (Wilson et al., 2012).

139 a.

b.

Figure 63. a) Basemap shows locations of seismic lines for jump correlation. b) Jump correlation for seismic line NBP0401-159 (left) west of Ross Island and seismic line MIS-1 (right) south of Ross Island. Colored lines denote interpreted seismic surfaces. Bold x markers indicate intersection of two seismic profiles. 140 Figure 64. Time-to-depth conversion velocity model. Black curve is derived from VSP data and the red curve is derived from whole core velocity measurements. The blue curve is derived from the black and red curves, and represents the time-to-depth conversion values utilized in this study.

141 Figure 65. Spatial extension of the Mt. Erebus (yellow polygon), Mt. Terror (red polygon), and Mt. Bird (blue polygon) volcanic loads. Basemap shows topography of McMurdo Sound onshore elevations denoted by yellow – white color ramp, bathymetric elevations denoted by blue color ramp).

142 Figure 66. 3D View of Ross Island volcanoes derived from digital elevation model.

143 Figure 67. Modelling results showing surface deformation (purple = greatest flexure) from Ross Island loading for increasing lithospheric Effective Elastic Thickness (Te).

144 Figure 68. Modelling result showing surface deformation (purple = greatest flexure) for Mt. Bird loading for a lithospheric Effective Elastic Thickness (Te) of 4 km.

145 Figure 69. Modelling result showing surface deformation (purple = greatest flexure) for Mt. Bird and Mt. Terror loading for a lithospheric Effective Elastic Thickness (Te) of 4 km.

146 Figure 70. Modelling result showing surface deformation (purple = greatest flexure) for Mt. Bird, Mt. Terror, and Mt. Erebus loading for a lithospheric Effective Elastic Thickness (Te) of 4 km.

147 Figure 71. Best fit model is represented by an overall minimum misfit for varying lithospheric Effective Elastic Thicknesses (Te). Horizontal axis is Te from 1-40 km. Vertical axis is RMS of misfit of dip angle of Rj surface on seismic lines VS dip angle of modeled flexural surface along same line (unit: degree). The best fit Te is 4 km.

148 Figure 72. Comparison of Rj surface depth (m) from seismic mapping (colored map with 100 m contour interval) with flexural surface (contours with 100 m contour interval) in response to Mt. Bird loading when lithospheric Effective Elastic Thicknesses Te = 4 km.

149 Figure 73. Comparison of Rk surface depth (m) from seismic mapping (colored map with 100 m contour interval) with flexural surface (contours with 100 m contour interval) in response to Mt. Erebus loading when lithospheric Effective Elastic Thicknesses Te = 4 km.

150 Figure 74. Interpretation of seismic line NBP0401-128, NBP0401-157, NBP0401-158. Colored lines denote interpreted seismic surfaces. Shaded blue regions are associated with seismic sequence 1 (SS1) and yellow regions are associated with seismic sequence 2 (SS2). Stippled pink regions designate volcanic aprons, and the shaded red region denotes a fissure ridge.

151 Figure 75. a) flexure model with island chain loading in map view (darker circles = migration of load through time). b) Stratigraphic pattern along the flank of island chain in response to island chain loading, assuming sediments filled to the same reference level after each loading (Watts and ten Brink, 1989).

152 Figure 76. 3D flexural modeling results from Aitken et al. (2012). Progressive flexural development of the southern McMurdo Sound basins with a lithospheric Effective Elastic Thickness (Te) of 3 km. The results have been cropped to show only Ross Island, relevant to this study; Aitken et al. (2012) included volcanoes to the south of Ross Island. 153 Figure 77. Results of 3D flexural modeling by Aitken et al. (2012). Upper panel shows transects through the 3D flexural model, showing the flexural surfaces obtained with a lithospheric Effective Elastic Thickness (Te) of between 0.5 km and 25 km. The gray shaded area indicates the range of flexure corresponding to Te between 2 and 5. Middle panel shows the predicted development of flexural accommodation space in southern McMurdo Sound with a Te of 3 km. Lower panel shows the location of C’-C profile across Ross Island.

154