INTERACTIONS BETWEEN TECTONISM AND MAGMATISM AT

SANTORINI: INSIGHTS FROM AN ACTIVE SOURCE SEISMIC EXPERIMENT

by

BEN HEATH

A DISSERTATION

Presented to the Department of Geological Sciences
and the Graduate School of the University of Oregon

in partial fulfillment of the requirements
for the degree of

Doctor of Philosophy

December 2019



DISSERTATION APPROVAL PAGE

Student: Ben Heath

Title: Interactions Between Tectonism and Magmatism at Santorini: Insights from
an Active Source Seismic Experiment

This dissertation has been accepted and approved in partial fulfillment of the
requirements for the Doctor of Philosophy degree in the Department of Geological
Sciences by:

Emilie E. E. Hooft Chair
Ilya Bindeman Core Member/Coordinator
Leif Karlstrom Core Member
Hank Childs Institutional Representative

and

Kate Mondloch Interim Vice Provost and Dean of the
Graduate School

Original approval signatures are on file with the University of Oregon Graduate
School.

Degree awarded December 2019

ii



© 2019 Ben Heath
All Rights Reserved

iii



DISSERTATION ABSTRACT

Ben Heath

Doctor of Philosophy

Department of Geological Sciences

December 2019

Title: Interactions Between Tectonism and Magmatism at Santorini: Insights from

an Active Source Seismic Experiment

In extensional volcanic arcs, tectonic and magmatic processes control

the evolution of eruptive centers and their surrounding regions. Faulting, which

increases crustal permeability, can focus magmatism and preferentially localize

eruptive edifices near intersecting fault zones. In turn, magmatic diking and

volcanic edifice growth/collapse can change both the regional and local stress/strain

in the crust altering the style and amount of faulting. The relative importance of

these magmatic and tectonic processes as well as how these processes coevolve are

still poorly known. Here we study magmatic and tectonic interactions at Santorini

Volcano, focusing on the localization of magmatism in the highly extended terrain

and the subsequent influence of this magmatism on the evolution of tectonic

activity. We use the dense PROTEUS active source seismic dataset, collected in

a 120 km x 45 km region around the volcano, to perform both P-wave anisotropic

traveltime tomography inversions and full waveform inversions, imaging the upper

crust down to ∼4 km depth. Integrating our results with the well-studied volcanic

and geologic history of Santorini, we show evidence for strong tectonic control on

the evolution of Santorini’s magmatic system. In Chapter II and V, this interaction

is recorded as i) the alignment of volcanic features parallel to tectonic features, ii)

the localization of volcanism within a tectonic basin, and iii) the strong influence of

tectono-magmatic lineaments on both regional volcanic evolution and localization

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of caldera collapse. In Chapter III we show that the magnitude of small-scale

local faulting is uncorrelated to regions of magmatic activity indicating that

magmatism is neither focused in areas of higher fracturing nor is accommodating

substantial extensional strain. Rather both magmatism and small-scale faulting

are strongly correlated with regional stress. In Chapter IV we hypothesize that the

broad distribution of deformation (>40 km wide) currently observed results from

magmatic activity. In contrast, a narrower episode of localized deformation (<20

km wide) preceded initiation of regional magmatism. This highlights the impact of

regional magmatism on tectonic evolution. This works shows that magmatic and

tectonic processes interact on a variety of temporal and spatial scales.

This dissertation includes both previously published and co-authored

material.

v



CURRICULUM VITAE

NAME OF AUTHOR: Ben Heath

GRADUATE AND UNDERGRADUATE SCHOOLS ATTENDED:

University of Oregon, Eugene, OR, USA
Northwestern University, Evanston, IL, USA

DEGREES AWARDED:

Doctor of Philosophy, Geological Sciences, 2019, University of Oregon
Master of Science, Geological Sciences, 2014, University of Oregon
Bachelor of Arts, Earth and Planetary Sciences, 2012, Northwestern

University

AREAS OF SPECIAL INTEREST:

Geophysics/Seismology
Volcanology
Tectonophysics

PROFESSIONAL EXPERIENCE:

Graduate Student, University of Oregon 2015-2019
ORISE Fellow, NETL-Albany 2015-2015
Graduate Student, University of Oregon 2012-2014

GRANTS, AWARDS AND HONORS:

Johnston Memorial Fellowship, 2012
Johnston Geology Fund, 2015, 2016, 2017
Weiser Memorial Scholarship, 2016, 2017
Outstanding Student Presentation Award, IAVCEI 2017
GSA Student Research Grant, 2017

PUBLICATIONS:

Heath, B. A., Hooft, E. E. E., Toomey, D. R., Papazachos, C. B., Nomikou,
P., Paulatto, M., Morgan, J.V., & Warner, M.R. (2019). Tectonism and
its relation to magmatism around Santorini volcano from upper crustal
P-wave velocity. Journal of Geophysical Research: Solid Earth DOI:
https://doi.org/10.1029/2019JB017699

vi



Hooft, E.E.E., Heath, B.A., Toomey, D.R., Paulatto, M., Papazachos, C.B.,
Nomikou, P., Morgan, J.V., & Warner, M.R. (2019). Seismic imaging
of Santorini: Subsurface constraints on caldera collapse and present-day
magma recharge. Earth and Planetary Science Letters, 514, 48-61. DOI:
https://doi.org/10.1016/j.epsl.2019.02.033

Hooft, E. E. E., Nomikou, P., Toomey, D. R., Lampridou, D., Getz, C.,
Christopoulou, M. E., O’Hara, D., Arnoux, G. M., Bodmer, M., Gray,
M., Heath, B. A., & VanderBeek, B. P. (2017). Backarc tectonism,
volcanism, and mass wasting shape seafloor morphology in the Santorini-
Christiana-Amorgos region of the Hellenic Volcanic Arc. Tectonophysics,
712, 396-414. DOI: https://doi.org/10.1016/j.tecto.2017.06.005

Heath, B. A., Hooft, E. E., & Toomey, D. R. (2018). Autocorrelation of the
seismic wavefield at Newberry Volcano: Reflections from the magmatic
and geothermal systems. Geophysical Research Letters, 45(5), 2311-2318.
DOI: https://doi.org/10.1002/2017GL076706

Mark-Moser, M. K., Rose, K. K., Schultz, J. D., Schultz, A., Heath, B.,
Urquhart, S., & Vincent, P. (2015). A Conceptual Geologic Model for
the Newberry Volcano EGS Site in Central Oregon: Constraining Heat
Capacity and Permeability Through Joint Interpretation of Multifaceted
Geophysical and Geological Data Sets (No. NETL-PUB-20152). NETL.

Heath, B. A., Hooft, E. E., Toomey, D. R., & Bezada, M. J. (2015). Imaging
the magmatic system of Newberry Volcano using joint active source and
teleseismic tomography. Geochemistry, Geophysics, Geosystems, 16(12),
4433-4448. DOI: https://doi.org/10.1002/2015GC006129

vii



ACKNOWLEDGEMENTS

I would first like to acknowledge my parents who have been supportive of

me throughout my graduate career. Without their help, I am not sure how I would

have finished this degree.

I would like to thank my advisor Emilie Hooft for both giving me an

opportunity to work on the PROTEUS dataset as well as for providing me the

space to develop and test my ideas. I would also like to thank my collaborators

during this experiment including: Emilie Hooft, Joanna Morgan, Paraskevi

Nomikou, Costas Papazachos, Michele Paulatto, Doug Toomey and Mike Warner.

Their feedback throughout this process has been invaluable. I am indebted to

Kajetan Chrapkiewicz and Brennah McVey for helpful conversations regarding

Santorini.

I would like to thank my committee which consists of Emilie Hooft, Ilya

Bindeman, Leif Karlstrom, and Hank Childs. I appreciate the support they

provided, their honest and open feedback, and their constructive criticism which

has helped me to become a better scientist.

My views on and understanding of magmatic systems has been shaped

by conversations with a wide range of people. The patience displayed by the

volcanology community in explaining concepts to a young scientist is greatly

appreciated.

I would like to thank my roommates, labmates and friends over the past few

years for their support. I feel honored to have gone through my graduate career

with so many bright people by my side.

Much of my research has used super-computers present on the University

of Oregon campus. I am constantly amazed by both my ability to make mistakes

viii



in the job submission process and the responsiveness of the ACISS/talapas high

performance computing team to my frantic cries for help. Without this team, I

would still be submitting “hello world” jobs to as many computer nodes as possible.

Finally, I would like to thank the Department of Earth Sciences (formerly

the Department of Geological Sciences or DoGS), especially the office staff, Marla,

Sandy and Dave.

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To my parents, thank you for always supporting me and helping me to become a

better person

x



TABLE OF CONTENTS

Chapter Page

I. INTRODUCTION . . . . . . . . . . . . . . . . . . . . . . . . 1

II. TECTONISM AND ITS RELATION TO MAGMATISM
AROUND SANTORINI VOLCANO FROM UPPER
CRUSTAL P-WAVE VELOCITY . . . . . . . . . . . . . . . . . . 5

2.1. Introduction . . . . . . . . . . . . . . . . . . . . . . . 5

2.2. Background . . . . . . . . . . . . . . . . . . . . . . . 6

2.2.1. Tectonic background . . . . . . . . . . . . . . . . . . 6

2.2.2. Volcanic Background . . . . . . . . . . . . . . . . . . 11

2.3. Experiment Geometry and Data Acquisition . . . . . . . . . 12

2.3.1. Seismic experiment . . . . . . . . . . . . . . . . . . . 12

2.3.2. Data return and data quality . . . . . . . . . . . . . . . 13

2.4. Data processing and tomographic inversion . . . . . . . . . . 14

2.4.1. OBS relocation . . . . . . . . . . . . . . . . . . . . . 14

2.4.2. Pg travel time data . . . . . . . . . . . . . . . . . . . 15

2.4.3. Tomographic Inversion . . . . . . . . . . . . . . . . . . 16

2.5. Results and Interpretation . . . . . . . . . . . . . . . . . 19

2.5.1. Upper crustal velocity variations . . . . . . . . . . . . . 19

2.5.2. Comparison of tomographic velocity model with
seismic reflection images . . . . . . . . . . . . . . . . . 20

2.5.3. Mapping geometry and thickness of sedimentary
basins from the tomography model . . . . . . . . . . . . 23

2.5.4. Western side of Santorini: Christiana Basin and
Christiana Ridge . . . . . . . . . . . . . . . . . . . . 24

2.5.5. Eastern side of Santorini: Anydros Basin, Anafi
Basin and Anydros Horst . . . . . . . . . . . . . . . . 27

2.5.6. Structures associated with Santorini . . . . . . . . . . . . 28

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Chapter Page

2.6. Discussion . . . . . . . . . . . . . . . . . . . . . . . . 29

2.6.1. Differences between the east and west sides of
Santorini and relationship to faulting . . . . . . . . . . . 29

2.6.2. Tectonic evolution of the basins and Christiana-
Santorini-Kolumbo magmatism: Localization of
magma in a proto-Anydros Basin . . . . . . . . . . . . . 32

2.6.3. Volcano-tectonic lineaments . . . . . . . . . . . . . . . 35

2.7. Conclusions . . . . . . . . . . . . . . . . . . . . . . . 36

2.8. Bridge . . . . . . . . . . . . . . . . . . . . . . . . . . 37

III. RELATIONSHIP BETWEEN FAULTING AND
MAGMATISM AROUND SANTORINI: SEISMIC
ANISOTROPY RESULTS FROM A REFRACTION EXPERIMENT . . 39

3.1. Introduction . . . . . . . . . . . . . . . . . . . . . . . 39

3.2. Background . . . . . . . . . . . . . . . . . . . . . . . 41

3.2.1. Geologic Background of the Aegean . . . . . . . . . . . . 41

3.2.2. Regional Volcanism and its Relation to Tectonics . . . . . . 43

3.2.3. Current State of Tectonic and Magmatic Activity . . . . . . 45

3.3. Methods . . . . . . . . . . . . . . . . . . . . . . . . . 46

3.3.1. Tomography . . . . . . . . . . . . . . . . . . . . . . 46

3.3.2. Forward Modeling . . . . . . . . . . . . . . . . . . . . 48

3.4. Results . . . . . . . . . . . . . . . . . . . . . . . . . 51

3.4.1. Tomographic Results . . . . . . . . . . . . . . . . . . 51

3.4.2. Forward Modeling Anisotropy Results . . . . . . . . . . . 57

3.5. Discussion . . . . . . . . . . . . . . . . . . . . . . . . 63

3.5.1. Source of Anisotropy on the eastern side of Santorini . . . . 63

3.5.2. Modeling the fraction of different aspect ratio fractures . . . 66

3.5.3. Magmatism and its relationship to
faulting/fracturing around Santorini - Local-
scale/Regional Scale Interactions . . . . . . . . . . . . . 68

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Chapter Page

3.5.4. Implications for Sources of Faulting and Fault
Damage Zones on the Eastern Side of Santorini . . . . . . . 73

3.5.5. Understanding the Regional Tectonic and
Magmatic Interactions . . . . . . . . . . . . . . . . . . 77

3.6. Conclusions . . . . . . . . . . . . . . . . . . . . . . . 79

3.7. Bridge . . . . . . . . . . . . . . . . . . . . . . . . . . 80

IV. TECTONISM AND THE INFLUENCE OF INHERITED
STRUCTURAL DEFORMATION AROUND SANTORINI
FROM P-WAVE TOMOGRAPHY . . . . . . . . . . . . . . . . . 81

4.1. Introduction . . . . . . . . . . . . . . . . . . . . . . . 81

4.2. Background . . . . . . . . . . . . . . . . . . . . . . . 83

4.3. Methods . . . . . . . . . . . . . . . . . . . . . . . . . 86

4.4. Results . . . . . . . . . . . . . . . . . . . . . . . . . 91

4.4.1. Overview of seismic anisotropy results from
travel-time tomography . . . . . . . . . . . . . . . . . 91

4.4.2. Overview of FWI results . . . . . . . . . . . . . . . . . 93

4.4.3. Comparison of FWI results to isotropic travel-
time tomography results . . . . . . . . . . . . . . . . . 97

4.4.4. Comparison of FWI to seismic reflection results . . . . . . . 99

4.5. Discussion . . . . . . . . . . . . . . . . . . . . . . . . 100

4.5.1. Seismic anisotropy and the record of ductile deformation . . . 100

4.5.2. Tomographic evidence of brittle behavior in the
Christiana Basin . . . . . . . . . . . . . . . . . . . . 105

4.5.3. Distribution of faulting within and outside the
basin proper . . . . . . . . . . . . . . . . . . . . . . 108

4.5.4. Ductile and brittle deformation: controls on the
localization of deformation, comparison between
east and west . . . . . . . . . . . . . . . . . . . . . . 114

4.6. Conclusions . . . . . . . . . . . . . . . . . . . . . . . 118

4.7. Bridge . . . . . . . . . . . . . . . . . . . . . . . . . . 121

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Chapter Page

V. SEISMIC IMAGING OF SANTORINI: SUBSURFACE
CONSTRAINTS ON CALDERA COLLAPSE AND
PRESENT-DAY MAGMA RECHARGE . . . . . . . . . . . . . . . 122

5.1. Introduction . . . . . . . . . . . . . . . . . . . . . . . 122

5.2. Santorini volcano . . . . . . . . . . . . . . . . . . . . . 124

5.3. Seismic experiment and data . . . . . . . . . . . . . . . . 129

5.4. Tomographic inversion and results . . . . . . . . . . . . . 130

5.4.1. Inversion . . . . . . . . . . . . . . . . . . . . . . . . 130

5.4.2. Synthetic tomography resolution tests . . . . . . . . . . . 133

5.4.3. Tomography results . . . . . . . . . . . . . . . . . . . 134

5.5. Calculation of physical properties . . . . . . . . . . . . . . 135

5.6. Discussion . . . . . . . . . . . . . . . . . . . . . . . . 139

5.6.1. Collapse plug filled with volcanic deposits . . . . . . . . . 143

5.6.2. Phreatomagmatic rock breakup . . . . . . . . . . . . . . 144

5.6.3. Multistage, nested caldera collapse . . . . . . . . . . . . 145

5.6.4. Present day magma recharge . . . . . . . . . . . . . . . 147

5.7. Conclusions . . . . . . . . . . . . . . . . . . . . . . . 150

VI. CONCLUSIONS . . . . . . . . . . . . . . . . . . . . . . . . . 152

APPENDICES

A. APPENDIX CHAPTER II: TECTONISM AND ITS
RELATION TO MAGMATISM AROUND SANTORINI
VOLCANO FROM UPPER CRUSTAL P-WAVE VELOCITY . . . . . 155

A.1. Description of Synthetics . . . . . . . . . . . . . . . . . 155

A.2. Attenuation of caldera centered data . . . . . . . . . . . . 168

A.3. First motions of signals . . . . . . . . . . . . . . . . . . 168

A.4. 1-D velocity profile comparison . . . . . . . . . . . . . . . 168

A.5. Investigations of roughness/smoothing values vs
RMS of travel time residuals . . . . . . . . . . . . . . . . 168

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Chapter Page

A.6. Dv analysis as function of penalty and smoothing . . . . . . . 168

A.7. Checkerboard test: 5x5x2 km checkers . . . . . . . . . . . . 168

A.8. Checkerboard test:5x5x1 km checkers . . . . . . . . . . . . 168

A.9. Checkerboard test: 3x3x1 km checkers . . . . . . . . . . . . 168

A.10. MCS comparison Supplement 1 . . . . . . . . . . . . . . . 168

A.11. MCS comparison Supplement 2 . . . . . . . . . . . . . . . 168

A.12. Contouring basin depth using various different velocities . . . . 168

A.13. Evidence for anisotropy in traveltime residuals . . . . . . . . 168

A.14. RMS of travel times for various different parameters . . . . . . 168

B. APPENDIX CHAPTER III: RELATIONSHIP BETWEEN
FAULTING AND MAGMATISM AROUND SANTORINI:
SEISMIC ANISOTROPY RESULTS FROM A
REFRACTION EXPERIMENT . . . . . . . . . . . . . . . . . . 169

B.1. Calculation of Resolution of Anisotropy . . . . . . . . . . . 169

B.2. Applicability of Anisotropy Assumptions . . . . . . . . . . . 169

C. APPENDIX CHAPTER IV: TECTONISM AND
THE INFLUENCE OF INHERITED STRUCTURAL
DEFORMATION AROUND SANTORINI FROM P-WAVE
TOMOGRAPHY . . . . . . . . . . . . . . . . . . . . . . . . . 172

C.1. Modeling Anisotropy . . . . . . . . . . . . . . . . . . . 172

C.2. Modeling Styles of Extension . . . . . . . . . . . . . . . . 176

D. APPENDIX CHAPTER V: SEISMIC IMAGING OF
SANTORINI: SUBSURFACE CONSTRAINTS ON
CALDERA COLLAPSE AND PRESENT-DAY MAGMA RECHARGE . 178

D.1. Temperature and pressure corrections for velocity . . . . . . . 178

D.2. Checkerboard resolution test . . . . . . . . . . . . . . . . 179

D.3. Cylindrical low-velocity anomaly resolution test . . . . . . . . 179

D.4. Comparison of shallow tomography velocity model
with seismic reflection profiles . . . . . . . . . . . . . . . 179

D.5. Comparison of shallow tomography velocity model
with seismic reflection profiles . . . . . . . . . . . . . . . 179

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Chapter Page

D.6. Interpretation of the low VP anomaly if porosity
were filled with partial melt . . . . . . . . . . . . . . . . 179

REFERENCES CITED . . . . . . . . . . . . . . . . . . . . . . . . 184

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LIST OF FIGURES

Figure Page

2.1. Tectonic setting of the Aegean . . . . . . . . . . . . . . . . . . 7

2.2. Experiment geometry and regional tectonic structures . . . . . . . . 10

2.3. Example of seismic data from two basins . . . . . . . . . . . . . . 14

2.4. Regional 1-D P-wave velocity-depth profiles . . . . . . . . . . . . 18

2.5. Tomographic velocity anomalies . . . . . . . . . . . . . . . . . . 19

2.6. Regional cross section locations . . . . . . . . . . . . . . . . . . 21

2.7. Comparison with seismic reflection profiles . . . . . . . . . . . . . 23

2.8. Analysis of regional basin depths . . . . . . . . . . . . . . . . . 25

2.9. Cross sections through the velocity model . . . . . . . . . . . . . 26

2.10. Comparison between caldera centered geology and
tomographic velocity anomalies . . . . . . . . . . . . . . . . . . 30

2.11. Model of tectonic evolution . . . . . . . . . . . . . . . . . . . . 33

3.1. Regional tectonic setting and experiment geometry . . . . . . . . . 42

3.2. Azimuthal distribution of travel-time residuals: evidence
for anisotropy . . . . . . . . . . . . . . . . . . . . . . . . . 52

3.3. Anisotropic structure of the eastern portion of the experiment . . . . 54

3.4. Anisotropic depth distribution . . . . . . . . . . . . . . . . . . 55

3.5. Anisotropy magnitudes across the experiment . . . . . . . . . . . . 57

3.6. Calculation of the effect of dikes on seismic velocity . . . . . . . . . 59

3.7. Calculation of the effect of fractures on seismic velocity . . . . . . . 61

3.8. Calculation of the effect of dipping faults on seismic velocity . . . . . 63

3.9. Calculation of combined fracture-induced anisotropy with
metamorphic-structure derived anisotropy . . . . . . . . . . . . . 64

3.10. Regional overview of fracture distribution . . . . . . . . . . . . . 68

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Figure Page

3.11. Hypotheses relating tectonism and magmatism . . . . . . . . . . . 71

3.12. Interpretations of regional structure . . . . . . . . . . . . . . . . 74

4.1. Recorded and modeled waveforms . . . . . . . . . . . . . . . . . 88

4.2. Phase residual comparison between final and starting model . . . . . 92

4.3. Comparison of ideal and actual coverage for FWI inversions . . . . . 93

4.4. Anisotropic travel-time tomography for various depths . . . . . . . . 94

4.5. FWI results for various depth slices . . . . . . . . . . . . . . . . 95

4.6. Cross sections through the FWI models . . . . . . . . . . . . . . 96

4.7. Overview of cross sections and regions used for averaging velocities . . 98

4.8. Comparison between FWI and traveltime tomography . . . . . . . . 99

4.9. Comparison of FWI to seismic reflection studies: example 1 . . . . . 101

4.10. Comparison of FWI to seismic reflection studies: example 2 . . . . . 102

4.11. Anisotropy resulting from metamorphic horst properties . . . . . . . 106

4.12. Anisotropy distribution for various geologically motivated blocks . . . 110

4.13. Orientation of anisotropy throughout the experiement . . . . . . . . 111

4.14. Hypothesis 1: Wide rifting from thermal perturbations . . . . . . . 117

4.15. Hypothesis 2: Accommodation zones and their effect on
distributed faulting . . . . . . . . . . . . . . . . . . . . . . . 119

4.16. Model of regional evolution . . . . . . . . . . . . . . . . . . . . 120

5.1. PROTEUS seismic experiment map . . . . . . . . . . . . . . . . 124

5.2. The geological and volcanic features of Santorini . . . . . . . . . . 127

5.3. Maps of P-wave seismic velocity at Santorini . . . . . . . . . . . . 131

5.4. Cross-sections of seismic velocity along profile A-A’ and B-B’ . . . . . 132

5.5. Predicted porosity for the northern caldera low-velocity
anomaly and the reference metamorphic profile. . . . . . . . . . . . 136

5.6. Conceptual model for the evolution of the LBA eruption
and formation of the low-velocity anomaly. . . . . . . . . . . . . . 142

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Figure Page

5.7. Comparison of the low-velocity anomaly with magma
recharge during the 2011-2012 unrest and predicted
lithostatic pressure at 4 km depth. . . . . . . . . . . . . . . . . 149

A.1. Attenuation of caldera centered data . . . . . . . . . . . . . . . 156

A.2. First motions of signals . . . . . . . . . . . . . . . . . . . . . 157

A.3. 1-D velocity profile comparison . . . . . . . . . . . . . . . . . . 158

A.4. Investigations of roughness/smoothing values vs RMS of
travel time residuals . . . . . . . . . . . . . . . . . . . . . . . 159

A.5. Dv analysis as function of penalty and smoothing . . . . . . . . . . 160

A.6. Checkerboard test: 5x5x2 km checkers . . . . . . . . . . . . . . . 161

A.7. Checkerboard test:5x5x1 km checkers . . . . . . . . . . . . . . . 162

A.8. Checkerboard test: 3x3x1 km checkers . . . . . . . . . . . . . . . 163

A.9. MCS comparison Supplement 1 . . . . . . . . . . . . . . . . . . 164

A.10. MCS comparison Supplement 2 . . . . . . . . . . . . . . . . . . 165

A.11. Contouring basin depth using various different velocities . . . . . . . 166

A.12. Evidence for anisotropy in traveltime residuals . . . . . . . . . . . 167

B.1. Approximating anisotropy using a weak elastic assumption . . . . . . 170

B.2. Predicted velocities and travel-times as a function of
azimuth for various different magnitudes of anisotropy . . . . . . . . 171

D.1. Checkerboard resolution test . . . . . . . . . . . . . . . . . . . 180

D.2. Cylindrical low-velocity anomaly resolution test . . . . . . . . . . . 181

D.3. Comparison of shallow tomography velocity model with
seismic reflection profiles. . . . . . . . . . . . . . . . . . . . . 181

D.4. Comparison of shallow tomography velocity model with
seismic reflection profiles. . . . . . . . . . . . . . . . . . . . . 182

D.5. Interpretation of the low VP anomaly if porosity were
filled with partial melt . . . . . . . . . . . . . . . . . . . . . . 183

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LIST OF TABLES

Table Page

4.1. Data Processing Steps . . . . . . . . . . . . . . . . . . . . . . . 87

4.2. Model Parameters . . . . . . . . . . . . . . . . . . . . . . . . . 89

4.3. Inversion Parameters . . . . . . . . . . . . . . . . . . . . . . . 90

A.1. RMS of travel times for various different parameters . . . . . . . . . . 168

C.1. Extensional Style Parameters . . . . . . . . . . . . . . . . . . . . 177

xx



CHAPTER I

INTRODUCTION

Globally, magmatic centers are found to strongly correlate with areas

of tectonic activity on a variety of spatial scales. Volcanic cones form linear

alignments that parallel the local kilometer-scale fault orientations, larger volcanic

edifices often lie near the intersection of multiple fault zones and at volcanic-arc

scales magmatic eruption volume is found to correlate with extension rate (e.g.

Muirhead et al., 2015; MacLeod and Sherrod, 1988; Acocella and Funiciello, 2010).

This interaction is commonly attributed to tectonic faults providing pathways

for the ascent of both magma and magmatic fluids as well as due to tectonic

stresses localizing magmatic activity (e.g. Kokkalis and Aydin, 2013; Magee et

al., 2013). Magmatism can then change tectonic activity by i) accommodating

extensional strain and modifying crustal stresses, which change fault density and

orientation and ii) weakening the crust through thermal and physical effects (e.g.

Parsons and Thompson, 1991; Parsons and Thompson, 1993; Casey et al., 2006;

Ebinger and Casey, 2001, Corti et al., 2013). However, despite strong evidence for

interaction, in many volcanic systems it is not clear whether tectonic or magmatic

stresses dominate regional evolution. This is further complicated because the

interaction between these two processes evolves through both time and space, as

the thermal and physical properties of the crust change through the addition of

magma/magmatic fluids.

While there is a substantial body of work investigating these tectonic and

magmatic interactions, studies have historically been limited by several issues.

Geologic studies, which are limited to the surface expression of faults and volcanic

deposits, have been biased against features that exist at depth and are limited to

rocks present in exposures. Geophysical methods, which are designed to probe the

1



current state of upper crustal continental magmatic systems, have been historically

of too poor resolution to sufficiently resolve local-scale tectonic and magmatic

interactions and have only been able to image regional faults. Due to the ease of

data collection in marine settings over land settings, the best studied volcanoes

from a seismic tomography perspective are preferentially located in the ocean. It is

precisely these oceanic regions that are inherently undersampled in geologic studies

due to limited outcrops and as a results are few volcanoes that are well studied

both geophysically and geologically. Therefore, few studies are able to combine

well-mapped geologic observations with dense geophysical constraints without

sacrificing resolution from either technique.

Here I probe upper crustal tectonic and magmatic interactions in the

marine setting of Santorini Volcano, Greece by combining both anisotropic P-

wave travel-time tomography and Full Waveform Inversions (FWI) with a well-

studied geologic history. Santorini, which sits on extended Aegean crust, is a prime

location to study tectonic and magmatic interactions because it has experienced

multiple episodes and styles of tectonic extension that were both accompanied

and unaccompanied by magmatism, it has a well-studied geologic and geochemical

record and history outlining episodes of both tectonism and magmatism and it is

located in a marine setting enabling the collection of dense active source seismic

experiment which provides unprecedented resolution of the continental upper crust.

By combining the geophysical results with geological and geochemical studies, we

are able to resolve deeper features, record subtle interactions between tectonic and

magmatic activity and build/test conceptual models of the timescales, length-scales

and styles of tectonic and magmatic interactions.

Chapter II provides a regional overview of faults and magmatic features

imaged using a dense marine-land seismic experiment. In this chapter, I conducted

2



isotropic P-wave travel-time tomography inversions to resolve seismic velocity and

ultimately resolve crustal structure beneath Santorini and the surrounding region.

This chapter validates the structures interpreted in later chapters, outlines broad

tectonic and magmatic interactions around Santorini including volcano-tectonic

lineaments that have focused regional volcanism and proposes a regional model for

tectonic evolution. This chapter was published with coauthors E. E. E. Hooft, D.

R. Toomey, C.B. Papazachos, P. Nomikou, M. Paulatto, J.V. Morgan, M.R. Warner

in Journal of Geophysical Research in a paper titled “Tectonism and its relation to

magmatism around Santorini volcano from upper crustal P-wave velocity."

Chapter III discusses anisotropic travel-time inversion results for the

Santorini-Kolumbo region, a region inclusive of both the large Santorini caldera

and the nearby, submerged Kolumbo volcano. Here I investigate the localization

of magmatism in this region from local and regional faults. Anisotropic travel

time results resolve local, small, kilometer-scale fault structure, providing the first

tomographic insight into small-scale interaction between magmatism and tectonic

activity at Santorini. Using a model that relates observed seismic velocity and

anisotropy to fault fraction and orientation, I show that regions of magmatism

have no correlation to the magnitude of local-scale faulting but do have a strong

relationship to regional-scale faulting, indicating that small scale faults are not

a first order control on magmatic activity. These results show that the regional

structure on the eastern side of Santorini is characterized by pervasive local faulting

and that regional extension has provided a favorable stress regime to allow the rise

of magmatic fluids.

Chapter IV investigates the largely amagmatic Christiana basin (western

side of Santorini) using anisotropic travel-time tomography and FWI, outlining

the style of extension in the Christiana basin in order to understand a transition

3



from narrow amagmatic extension within the basin to broader magmatic extension

observed regionally. I investigate sources of seismic anisotropy here and show

that it results from a variety of different processes including various oriented

faults from differing extensional episodes as well as ductile-deformation-derived

metamorphic stretching lineations that record prior Oligo-Miocene extension. Using

FWI, I outline the regional tectonic fault structure of the Christiana basin and

investigate localization of faulting within the basin proper, inferring changes in

the orientation of faulting through time. The differences between older amagmatic

localized faulting within the Christiana basin and younger, distributed NE-SW

oriented faults associated with magmatism (discussed in Chapters II and III) is

hypothesized to be caused by thermal perturbations to the crust resulting from

initiation of magmatism. I provide an overview of different extensional modes and

their evolution through time.

Chapter V consists of a focused P-wave travel-time tomography study of

the structure of the Santorini caldera, highlighting the localization of vents around

a newly discovered region of isolated caldera collapse which in turn is localized

between two long-lived tectono-magmatic structures. This chapter, titled “Seismic

imaging of Santorini: Subsurface constraints on caldera collapse and present-day

magma recharge" was written by EEE Hooft, and coauthored by Heath, B.A.,

Toomey, D.R., Paulatto, M., Papazachos, C.B., Nomikou, P., Morgan, J.V. and

Warner, M.R. and is published in the journal Earth and Planetary Sciences. This

work is included in this dissertation as B. A. Heath collected all the travel-times

used in this project, conducted all tomographic inversions and contributed to the

development of ideas presented.

4



CHAPTER II

TECTONISM AND ITS RELATION TO MAGMATISM AROUND SANTORINI

VOLCANO FROM UPPER CRUSTAL P-WAVE VELOCITY

This chapter was published in Journal of Geophysical Research: Solid Earth,

under the title “Tectonism and its relation to magmatism around Santorini volcano

from upper crustal P-wave velocity.” This publication was co-authored with E. E.

E. Hooft, D. R. Toomey, C.B. Papazachos, P. Nomikou, M. Paulatto, J.V. Morgan,

M.R. Warner. Ben Heath picked the data used in this chapter, conducted all the

inversions, wrote the paper and drafted all figures. All co-authors contributed to

discussion of ideas and provided extensive edits to the document.

2.1 Introduction

In extensional volcanic arcs, the crust is often composed of a patchwork of

interacting faults that control the localization of magmatism and development of

sedimentary basins. Analogue models have shown that magma may be localized

within basins in different extensional regimes (Corti et al., 2003). In turn magma

bodies can alter the stress state of the crust. In addition to weakening the crust

through thermal effects, magmatism can facilitate the formation of low angle faults

(Parsons & Thompson, 1993), inhibit the formation of through-going normal faults

(e.g., Faulds & Varga, 1998), and accommodate extensional strain.

Santorini is one of the most active volcanoes in the Hellenic Volcanic arc

(e.g., Le Pichon & Angelier, 1979) (Figure 2.1), resulting from the subduction

of the African plate under the Aegean microplate. Regional extension, since the

Oligocene-Miocene, has been accommodated both on older E-W oriented faults

and on younger NE-SW oriented normal faults that remain tectonically active to

present day. As a consequence of this extension, the crust has been thinned and

5



multiple fault systems of differing orientations intersect Santorini (Figures 2.1

and 2.2). The link between tectonics and magmatism across Santorini, however,

is currently not well constrained, in part because many of these features are buried

under sedimentary and volcanic deposits.

In this study we use seismic data, acquired in the marine-land active-

source PROTEUS (Plumbing Reservoirs Of The Earth Under Santorini) seismic

experiment, to obtain a tomographic P-wave velocity model of the upper-crustal

structure across Santorini volcano and the surrounding region. Seismic velocity is

sensitive to fracturing, porosity, and composition and, here, we use it to delineate

the structure and evolution of faults, extensional basins, as well as volcanic features

in the Santorini region. We draw on our tomographic images, and results from

previous seismic reflection and other studies, to delineate basins and faults. The

combined datasets, as well as results from prior studies, are used to investigate

differences between the upper crust to the west and east of Santorini, build a

geologic model of fault evolution, and explore the relationship between tectonism

and magmatism.

2.2 Background

2.2.1 Tectonic background

The Aegean has undergone multiple episodes of extension from at least the

Oligocene-Miocene to present that were driven by slab roll-back of the Hellenic

subduction zone (Le Pichon & Angelier, 1979), westward extrusion of Anatolia

(McKenzie, 1972), and gravitational collapse of continental crust (McKenzie, 1972).

These extensional events have led to thinned crust (Makris, 1976, 1978), exhumed

and rotated metamorphic blocks (e.g. Walcott & White, 1998), and localized

6



20˚ 22˚ 24˚ 26˚ 28˚ 30˚
32˚

34˚

36˚

38˚

40˚

42˚

Methana

Milos
Santorini

Nisyros&Kos

~30 m
m

/yr

Aegean

Greece

Turkey

Africa

Hellenic Trough

North Anatolia Fault

Cyclades

Anatolia

Crete

Cretan Basin

Eastern Aegean

volcanic arc

volcano

study area

approximate 
gps velocity

Figure 2.1. Tectonic setting of the Aegean. Topographic/bathymetric map of
the broader Aegean/Hellenic subduction zone (bold grey line with triangles) where
the African plate subducts under the Aegean microplate showing the volcanic arc
(red region) and main volcanic centers (stars and bold names). The study area is
indicated with a dashed box. Faults in the overriding Aegean microplate (black
lines, after Caputo et al., 2012) result from subduction rollback and extrusion
along the North Anatolia Fault. The Santorini study area is located between the
Cyclades to the northwest and the Cretan Basin to the south, within an actively
faulted region. The GPS-derived velocity vectors relative to Eurasia (black arrows)
reflect differential motion between regions east and west of Santorini (Reilinger et
al., 2010).

7



deformation along zones of crustal weakness (Jolivet et al., 2013 and citations

therein), creating a variety of deforming terrains.

Geological studies show that the Cyclades (Figure 2.1) were extended,

thinned, and rotated throughout the Miocene, but have behaved as a single

rotating block since the late Miocene-early Pliocene (Walcott & White, 1998) and

currently exhibit little internal deformation. Santorini and the surrounding region

lies at the southeastern margin of this Cycladic block (Le Pichon & Kreemer, 2010)

(Figure 2.1).

During the Miocene to Pliocene, E-W-striking normal faults and related

basins formed under ~N-S extension (Anastasakis & Piper, 2005; Piper et al.,

2007). One of these basins is the Christiana basin in the western part of the study

area. The Christiana basin has been mapped using seismic reflection imaging

(Tsampouraki-Kraounaki & Sakellariou 2018, Piper et al., 2007 and citations

therein) and is thought to have been truncated against N-S-striking transfer faults

(Piper & Perissoratis, 2003). This E-W-striking basin is filled with Messinian

and younger sediments (Piper & Perissoratis, 2003, Tsampouraki-Kraounaki &

Sakellariou 2018). The main faults bounding the Christiana basin ceased activity

~1 Ma during the Pleistocene (Piper and Perissoratis, 2003) but some activity

continues to present day as evidenced by small-offset, active faults observed in

seismic reflection data (Tsampouraki-Kraounaki & Sakellariou, 2018).

In the latest Pliocene or early Pleistocene, the regional stress direction

rotated resulting in NE-SW trending faults and subsequent basin inversion

(Piper and Perissoratis, 2003; Piper et al., 2007) as well as the formation of new

basins such as the Anydros and Anafi Basins on the eastern side of our study

area (e.g., Hübscher et al., 2015; Nomikou et al., 2016). Following the change in

8



faulting orientation, volcanism in the region began (Perissoratis, 1995), with initial

volcanism sourced at Christiana and Akrotiri (SW Santorini) (Piper et al., 2007).

Faults striking NE-SW continue to be active to present day (e.g.,

Dimitriadis et al., 2009, Bohnhoff et al., 2006) and are thought to be predominately

extensional to trans-tensional (Hubscher et al., 2015, Hooft et al., 2017, Nomikou et

al., 2016, Nomikou et al., 2018), but a component of strike-slip behavior is observed

(Sakellariou et al., 2010). The NE-SW-striking Santorini-Amorgos fault zone (e.g.

Stiros et al., 1994), which includes the Anydros Region on the eastern side of the

study area (Figure 2.2), is one of the most active in the Aegean and regionally is

the margin between the seismically quiet Cycladic block and the diffuse seismicity

observed in the eastern Aegean (Bohnhoff et al., 2006). This is consistent with GPS

studies showing that the eastern Aegean moves to the SE relative to the Cylcadic

block (McClusky et al., 2000; Reilinger et al., 2010).

Metamorphosed sediments of the Attico-Cycladic complex, which were

exhumed during Miocene and younger extension (Piper et al., 2007; Piper &

Perissoratis, 2003; Xypolias et al., 2010), highlight the role of tectonics in the

evolution of the southern Aegean. These exhumed rocks compose most of the

regional islands (e.g. Ios, Amorgos) ( Lister et al., 1984; Forster et al., 1999),

including a portion of SW Santorini itself (Heiken & Mccoy, 1984). Moreover, these

metamorphic fault blocks often extend as horsts into the Aegean sea, as seen in sea

bathymetry and seismic reflection profile images (Perissoratis, 1995; Hübscher et

al., 2015; Nomikou et al., 2016; 2018; Hooft et al., 2017). Metamorphic basement

lithics, found in the erupted ignimbrite products from Santorini, reveal that the

metamorphic complex must underlie Santorini’s caldera (Druitt, 2014) and it is

thought that Santorini sits on a NE-SW oriented fracture network (Budetta et al.,

1984).

9



25˚00'

25˚00'

25˚15'

25˚15'

25˚30'

25˚30'

25˚45'

25˚45'

26˚00'

26˚00'

36˚15' 36˚15'

36˚30' 36˚30'

36˚45' 36˚45'

25˚00'

25˚00'

25˚15'

25˚15'

25˚30'

25˚30'

25˚45'

25˚45'

26˚00'

26˚00'

36˚15' 36˚15'

36˚30' 36˚30'

36˚45' 36˚45'

−1400 −1200 −1000 −800 −600 −400 −200 0 200 400

Kameni Islands

Amorgos B
asin

Christiana Basin (CB)

Anydro
s H

orst
 (A

H) 

Ana� Basin (AFB)

Christiana
Akrotiri

Anydros B
asin

 (A
B)

Santorini
Ana�

Ios

Ios f
ault z

one (IF
)

Anydros

Amorgos

 Sea of Crete

Ak
ro

tir
i 

fa
ul

t

Santorini Ana� Basin (SAB)

Ana� Amorgos Basin (AAB)

Kolumbo Volca
nic 

Chain

KL1

KL2

KL1 = Kolumbo Line
KL2 = Kameni Line
NC  = Northern Caldera
SC   = Southern Caldera

Christiana Ridge (CR)

NC

SC

Kolumbo

Any
dro

s F
au

lt 
(A

F)

Christiana 
fault

148

105

a)

b)

Cretan Basin 

x
y

Figure 2.2. The Santorini region showing the PROTEUS active seismic
experiment and the main morphotectonic features. (a) Topographic/bathymetric
map of Santorini and the surrounding islands showing seafloor and island seismic
stations (yellow circles are stations used in inversion, grey circles are stations that
were not used) and the seismic shot profiles (small red circles). Data from seismic
lines recorded on stations 148 and 105 (blue/purple) are shown in Figure 2.3. (b)
Topographic/bathymetric map with the main basins, horsts/ridges, and faults (red
lines, from Hooft et al., 2017). Abreviations used are: CB-Christiana Basin, CR -
Christiana Ridge, IF - Ios Fault, Anydros Horst, AB - Anydros Basin, AFB -Anafi
Basin, SAB - Santorini-Anafi Basin, AAB - Amorgos-Anafi Basin. Kolumbo and
Kameni lines (KL1 and KL2, respectively) are after Pfeiffer (2001).

10



2.2.2 Volcanic Background

Regional volcanism has been localized in three areas: Christiana Island,

Santorini volcano and the Kolumbo volcanic chain, arranged roughly in a NE-SW

direction (Nomikou et al., 2013) (Figure 2.2). The oldest of the regional volcanic

deposits results from Christiana volcanism (active during late Pliocene) and from

early volcanic centers at the SW edge of Santorini near Akrotiri (~0.6 Ma) (Piper

et al., 2007). Volcanism at Santorini itself began roughly 650 ka (Druitt et al.,

1999), with explosive volcanism beginning around 360 ka. At least 4 major caldera-

forming eruptions have occurred at Santorini (Druitt, 2014), the most recent of

which, the 30-86 km3 (dense rock equivalent) Late Bronze Age (LBA) "Minoan"

eruption, occurred 3.4 ka (Druitt, 2014 and citations therein; Johnston et al., 2014;

Friedrich et al., 2006).

Post-LBA effusive eruptive vents are located along the Kameni line, a NE-

SW-trending line of vents in the caldera (Nomikou et al., 2014; Pyle & Elliott,

2006) that results from either a deep seated fault focusing magmatism, a caldera

ring fault, or a region of diking (Konstantinou et al., 2013; Newman et al., 2012;

Papadimitriou et al., 2015; Saltogianni et al., 2014; Tassi et al., 2013). The

Kolumbo line, a similarly oriented lineament north of the Kameni line (see Figure

2.2), is observed as a linear alignment of older volcanic centers in the NE caldera

and is thought to strike toward the Kolumbo volcanic chain NE of Santorini. The

Kolumbo volcanic chain is a NE-striking alignment of submarine seamounts (e.g.,

Hooft et al., 2017; Nomikou et al., 2012; 2013), the largest of which is Kolumbo

Seamount (Hübscher et al., 2015; Nomikou et al., 2016).

The interaction between magmatism at Santorini and Kolumbo is still

debated. Petrologic studies indicate that their respective crustal magmatic systems

are not linked (Klaver et al., 2016), though earthquake tomography suggests a

11



possible connection below 5 km depth (Dimitriadis et al., 2010). The NE-SW strike

of regional volcanism, combined with NE-SW striking fault zones, has led other

researchers to suggest an interaction between magmatism and tectonic stresses

(e.g., Dimitriadis et al., 2009; Feuillet, 2013). Similar NE-SW trending faults are

also present at other large Aegean volcanoes (Papazachos & Panagiotopoulos, 1993;

Nomikou & Papanikolaou, 2011).

2.3 Experiment Geometry and Data Acquisition

2.3.1 Seismic experiment

In November and December of 2015, a three-dimensional, active-source

seismic tomography experiment was conducted in the broader Santorini volcano

area of the southern Aegean Sea. The experiment covered an area roughly 120

km x 45 km centered on Santorini, with the goal of imaging the crustal magmatic

plumbing system beneath the volcano (Figure 2.2). Data were collected on 91

ocean bottom seismometers (OBSs) and 65 land seismometers distributed around

Santorini’s caldera, as well as on Anafi, Christiana, and Anydros islands. The

OBSs included 30 Woods Hole Oceanographic Institution (WHOI) instruments

(4.5 Hz Geospace 3-axis geophone) and 61 Scripps Institute of Oceanography (SIO)

instruments (4.5 Hz Sercel L-28 3-component geophone). Both OBS types also

had a High Tech HTI-90-U hydrophone. The land seismometers included 60 Mark

1-Hz geophones from the German Research Center for Geosciences Geophysical

Instrument Pool in Potsdam and 5 CMG-40T and Trillium compact (120s)

seismometers from the Aristotle University of Thessaloniki. All OBS instruments

had a 200-Hz sampling rate and all land stations used a 100-Hz sampling rate.

There were over 14,300 active-source marine shots using a 6600 in3 36-component

airgun array towed at 12 m depth. Average shot spacing was ~150 m along ENE-

12



WSW-oriented shot lines spaced at 1-2 km intervals and included shot-receiver

offsets up to ~115 km. Additional azimuthal coverage was achieved using a lower

shot-density NW of Santorini and stations located on the neighboring Anafi island

(Figure 2.2a).

2.3.2 Data return and data quality

The OBS data return was sufficient with 1 OBS lost and an additional 10

OBSs with noisy and/or poor-quality records on both the hydrophone and vertical

channels. A sporadic ~6 Hz ringing noise was observed on several SIO stations

that did not correlate to any external processes such as wind, lightning, boat

traffic, etc. This noise was of variable amplitude and occurred randomly on all 4

channels. Figure 2.3 provides an example of good quality OBS data for two stations

located in the Christiana and Anydros basins. Differences between the basins are

clearly observed, with large, rather impulsive first arrivals with short coda in the

Christiana basin and more chaotic arrivals with a "ringy" character and longer coda

in the Anydros basin.

Santorini land stations had variable quality data. Those located on

metamorphic rocks and exposed bedrock had great return; in contrast those

situated on volcanic deposits were noisy and difficult to pick. Clear impulsive

arrivals were seen on Anafi stations, as all these stations were installed on

metamorphic basement, providing good longer-range azimuthal coverage to the

experiment. However, the Anafi stations were located outside the velocity model

used in this paper and thus picks for these stations were not included.

For upper crustal Pg waves crossing the caldera area, arrivals were highly

attenuated and travel times for these stations were difficult to pick due to emergent

waveforms (Figure A.1). For seismic waveforms that did not travel within the

13



0 10 20 30 40

3

2

1

0

0 10 20 30 40

3

2

1

0

tim
e 

- x
/6

   
(s

)

x-distance (km)

3 
km

10
 k

m

20
 k

m

10
 k

m

20
 k

m

5 
km

10
 k

m

15
 k

m

20
 k

m

25
 k

m

5 
km

10
 k

m

15
 k

m

W E

W E

Christiana Basin

Anydros Basin

Station 148: Hydrophone

Station 105: Hydrophone

Range (km)

a)

b)

tim
e 

- x
/6

   
(s

)

Figure 2.3. Seismic data recorded in basins east and west of Santorini. Seismic
record sections (hydrophone) for shot lines in the Christiana (top) and the Anydros
Basins (bottom) shown with a reduction velocity of 6 km/s. Shots in the Christiana
Basin are characterized by a large impulsive first motion in contrast to the more
’ringy’ character of those in the Anydros Basin. There is good agreement between
travel time picks (red) and predicted travel times through our preferred model
(blue). Regions shaded grey correspond to near shot ranges (0-6 km) which were
excluded from the inversion (see text for details).

footprint of the volcano, clear, impulsive arrivals were observed, indicating that

any attenuative structure was located beneath the caldera.

2.4 Data processing and tomographic inversion

2.4.1 OBS relocation

The OBSs were relocated on the seafloor by minimizing the misfit between

predicted and observed acoustic water-wave arrival times (Creager & Dorman,

1982) for short range arrivals (0-2 km) (longer ranges were included when

14



necessary). A constant water velocity of 1.52 km/s (determined from expendable

bathythermographs) was used and a pick error of 5 ms was estimated. Station

locations were fixed to the seafloor using the bathymetric map from Hooft et al.

(2017). OBS depths ranged from ~30 m to ~890 m, with typical depth uncertainties

of 10 m and horizontal uncertainties less than 4 m. Using the precision of the

water-wave arrivals (most stations were fit to water-wave travel-time RMS values of

~5 ms), we were able to identify and correct several errors in the dataset including

origin time offsets and mis-located shots. In addition, a non-linear response at short

ranges (typically < 0.8 km) was observed on the hydrophone for many instruments,

especially shallow stations. This effect did not influence picking of the first motion

of water-waves, although it did affect the resulting waveform shape (Figure A.2).

Finally, we were able to identify and correct for a linear drift in the OBS clock for

one instrument that had not been accounted for in the initial data processing.

2.4.2 Pg travel time data

We picked over 200,000 Pg first arrivals using two approaches. Initial picking

was conducted using an autopicker from the open-source opendTect software

(https://www.dgbes.com/). Several tens of thousands of arrivals with high signal to

noise ratio (SNR) traces (e.g. short ranges, predominately less than 20 km range)

were collected using this approach. The picking error for these arrivals was assigned

a value of 10 ms, following visual inspection of the automatic picks. After this

first step, arrivals were picked manually. Errors were visually assigned during the

picking process and ranged from 5 ms (high SNR, exceptional quality data) to 30

ms (low SNR); the median error was 10 ms with standard deviation of 13 ms. The

difference in assigned errors resulted from both variability in the noise level, as well

as differences in waveform shape and arrival time due to surface and subsurface

15



complexity. Data were picked on the hydrophone, vertical, and/or a scaled linear

stack of the hydrophone and vertical channels. All picks (manual and automatic)

were made on the first negative to positive zero crossing (using the polarity

convention of the hydrophone channel, Figure 2.3) as it is the easiest portion of

the first arrival pulse to pick. This zero-crossing followed a small negative first

motion that was only easily observed in the highest quality data. All picks were

made on waveforms filtered with a 4th order causal Butterworth filter of 5-25 Hz,

for shots with distance ranges between 0 and 65 km. Picking resulted in a high-

quality dataset for distances between 4 and 30 km at most stations.

2.4.3 Tomographic Inversion

We inverted the Pg travel time first arrivals using the approach of Toomey

et al. (1994), minimizing the squared residual between observed and predicted

travel times while also penalizing against both the magnitude and roughness

of model perturbations (roughness measured using Laplacian smoothing). The

slowness model was defined on a 120 km x 45 km x 12 km (x, y, and z dimensions)

rotated rectangular grid for the forward problem, with a grid spacing of 200 m for

both horizontal (x and y) and vertical (z) dimensions, and bathymetry reflected

by vertical shearing of the grid (Toomey et al., 1994). The grid was rotated 25.5o

counter-clockwise from north, to align the x-axis with the shot-line orientation

(Figure 2.2a). Travel times through the crust were calculated using the graph

theory approach of Moser (1991), with the water-wave segment of the travel-time

(shot to seafloor) calculated on a grid with a 50 m x 50 m horizontal spacing, using

a constant water-wave velocity of 1.52 km/s. For the linearized model inversion,

a perturbational grid of 400 m x 400 m x 200 m in the horizontal and vertical

directions, respectively, was used, inverting for slowness perturbations. Updated

16



slowness values were then linearly interpolated onto the forward problem grid.

Iterating on this linearized inversion strategy allowed for an accurate approximation

to the non-linear travel time tomographic problem (see Toomey et al., 1994 for

additional details) A number of inversions were conducted to create a smooth 3-D

VP starting model, so that large velocity contrasts resulting from the metamorphic

horsts and sediment-filled grabens were included in the starting model. To first

obtain the best 1D velocity model for the region (Figure 2.4), a 3D inversion was

conducted using a 1-D VP starting model derived from a combination of prior

gravity, seismic refraction, and receiver function studies (Bohnhoff et al., 2006)

(blue line in Figure A.3). The 3-D output velocity model was averaged at each

depth and the new 1-D model was used as an updated starting model. This process

was repeated and allowed for the migration to an optimized regional 1-D velocity

model (black line Figures 2.4 and A.3). To obtain the smooth 3-D VP starting

model, we used this regional 1-D velocity model and inverted for 3-D isotropic

velocity variations, which were then spatially smoothed with a median filter of

5 km by 5 km by 2 km in the horizontal (x and y) and vertical (z) directions,

respectively. For the final inversion we used horizontal and vertical smoothing

parameters of 200 and 100, respectively, and penalized model perturbations

relative to the previous model iteration using a penalty of 1 (see Toomey et al.,

1994 for more detail). Each inversion consisted of 5 model iterations, ensuring the

convergence of the RMS travel time misfit (Figure A.4). We conducted tens of trial

inversions and the main model features presented in this paper were insensitive

to reasonable variations of the inversion parameters (Table A.1, Figures A.4 &

A.5). Ranges shorter than 6 km were poorly fit in the initial inversions, a result

of the lack of short-range travel-time picks and a large grid spacing (200 x 200 x

200 m) relative to the ranges. Picks associated with these ranges were therefore

17



1 2 3 4 5 6 7
Velocity (km/s)

0

1

2

3

4

D
ep

th
 (k

m
)

10%

0%

5%

Intrusions and 
metamorphics

Probability 

Figure 2.4. 1-D P-wave velocity-depth profiles. Averaged 1-D velocity depth
profile of four selected areas around Santorini (outlined in subfigure), compared to
the average regional velocity profile (black line). Basin fill is assumed for velocities
less than 4 km/s and metamorphic rocks and intrusives for velocities faster than 4
km/s - this corresponds to a steep gradient in the velocity-depth curve. The light
blue curve highlights the P-wave velocity for the Anydros metamorphic horst (Fig.
2.2), showing faster than average velocities of the metamorphic rocks at shallow
depths (0-2 km). At > 2 km depth, areas on the western side of Santorini (red
and magenta) are systematically faster than those on the eastern side (dark blue
and light blue). Background grey-scale shows the probability density function
distribution of all 1-D velocities of the tomography model.

removed from the inversions. The model presented here (Figure 2.5) was fit to an

RMS travel-time misfit of 15 msec, which corresponds to a χ2 of 2.2. To ensure the

accurate recovery of structures, initial inversions only used the higher quality, often

shorter-range data (< 15 km). By predicting expected (theoretical) arrival times

through this initial tomographic model, we were able to identify mis-picked arrivals

and facilitate the manual picking of noisy first arrivals, which were skipped in the

initial picking effort. Using this approach, we extended the picked dataset from <

15 km (high signal-to-noise) to > 30 km (often lower signal-to-noise).

18



b)

2 km

c)

AH

CR

CB

AB

AFB

IF

SAB

AAB

10 km

0 km

a)

AH

CR

CB

AB

AFB

IF

SAB

AAB

10 km

3 km

d)

AH

CR

CB

AB

AFB

IF

SAB

AAB

10 km

1 km 

10 km

AH

CR

CB

AB

AFB

IF

SAB

AAB

b)

dv km/s

Figure 2.5. Velocity anomalies relative to the 1-D average velocity-depth profile
at 0, 1, 2, and 3 km depth. Background is illuminated bathymetric map from Hooft
et al. (2017). Grey lines are faults defined using the seismic tomography images.
These faults are colored black on the depth slice on which they were defined.
Dashed black lines are large faults outside the region of resolution. Abbreviations
as in Figure 2.2b.

2.5 Results and Interpretation

2.5.1 Upper crustal velocity variations

The tomographic model shows substantial P-wave velocity variations in the

upper crust of the study area, with lateral differences in velocity exceeding 3 km/s

near the surface (Figure 2.5). In the top-most kilometer, regions of anomalously

high (+1.5 km/s) and low (-1.5 km/s) relative velocity correspond to mapped

metamorphic basement and sedimentary basins, respectively (Figures 2.2 & 2.5a-

b). At depths greater than 2 km, the velocity west of Santorini is higher and

19



spatially more uniform compared to the eastern side of the Santorini volcano, where

velocities are both lower and more variable (Figures 2.4 & 2.5b). At 3 km depth,

the overall lateral velocity variability across the study area is still ~2 km/s (±

1 km/s) (Figure 2.5d). The reliability of the recovered longer wavelength lateral

velocity variations is validated by checkerboard resolution tests (Text A.1, Figures

A.6, A.7 & A.8), which show that features with length scales of 5 km and greater

are well recovered throughout the model, while features with length scales of 3 km

are well recovered within the higher-velocity areas, mainly metamorphic basement.

Rapid velocity changes between the low-velocity sedimentary basins and

higher velocity metamorphic basement rocks are interpreted as faults. Near-

surface faulting, delineated by sharp velocity changes observed in the 0- and 1-km-

depth slices (heavy black lines in Figure 2.5a&b) is in good agreement with faults

observed in the topographic-bathymetric map (Figure 2.2). At the depth of 2 and 3

km, we interpret similar sharp spatial velocity gradients as faults within or beneath

the basins (heavy black lines in Figure 2.5c-d).

2.5.2 Comparison of tomographic velocity model with seismic reflection images

In general, first order geotectonic features, such as large-scale sedimentary

basins and their bounding faults, are well resolved in both tomographic and

prior seismic reflection datasets and show similar geometries and structures.

Figure 2.6 shows the location of 7 cross-sections where detailed model evaluation

and interpretation were performed, of which profiles A-A’, B-B’ and C-C’ have

associated seismic reflection results (Figure 2.7).

Figure 2.7a compares the seismic tomography cross-section to a multi-

channel seismic (MCS) profile that crosses the Anydros Basin, Anydros Horst,

and Anafi-Amorgos Basin from NW to SE (Nomikou et al., 2018) (Figure 2.6).

20



D 

D’

E

E’

F

F’

G
G’

A

A’

B B’

C

C’

Figure 2.6. Map showing the location of seismic cross-sections. A-A’, B-B’ and C-
C’ (purple lines) indicate the locations of the three seismic reflection profiles (from
prior studies) and corresponding velocity cross-sections (this study) that are shown
in Figure 2.7. D-D’, E-E’, F-F’ and G-G’ (red lines) show the locations of the four
interpreted tomography cross-sections through regional basins shown in Figure 2.9.
Tomographically defined faults (black lines) are after Figure 2.5. Orange regions
mark areas of diffuse micro-seismicity (Bohnhoff et al., 2006; Brüstle et al, 2014).

The transition across a large-offset normal fault into the Anafi-Amorgos basin

accurately coincides with an abrupt change from high to low velocities in the

tomographic image. A smaller basin bounded by secondary faults is similarly

expressed tomographically in the Anydros Horst footwall. The transition from

sediment to the underlying basement of the Anafi-Amorgos basin corresponds

approximately to the 4 km/s velocity contour. Additional comparisons with the

MCS results (Nomikou et al., 2016) in the vicinity of Kolumbo submarine volcano

are presented in Figures A.9 and A.10.

A seismic tomography comparison to a W-to-E seismic reflection profile

within the Christiana Basin (Tsampouraki-Kraounaki & Sakellariou 2018) is

presented in Figure 2.7b. In the reflection image, the basin is filled by sediments

(Units 1 -5, Figure 2.7b) intercalated with pyroclastic flows (Roman numerals),

21



with Unit 6 proposed to be Messinian evaporites, a late Miocene marker

(Tsampouraki-Kraounaki & Sakellariou 2018). In the center of our tomographic

velocity profile, elevated seismic velocities relative to the surrounding sediments

(3 - 4 km/s versus 1.5 - 3 km/s) correlate with an updoming of the Messinian

evaporites, an observation consistent with research showing that evaporites are

seismically faster than sediments and pyrolastic flows (e.g., Zong et al., 2017).

Small-offset faults within the sediments on the eastern end of the reflection profile

correlate with underlying variations in basement structure in the tomography

images.

Figure 2.7c shows a S-N seismic reflection profile through the Christiana

Basin (Tsampouraki-Kraounaki & Sakellariou 2018) and its comparison with

tomographic results. Generally, the main basin is seismically slow, whereas the

basement rock (Unit 7 in top figure) is seismically fast (~3.5 to 5.5 km/s). In the

northern portion of the basin the dip of the margin is well recovered. An old buried

basin margin on the southern edge of the Christiana basin, thought to have become

inactive ~1 Ma (Tsampouraki-Kraounaki et al. 2018), is also tomographically

identified as seismically faster than the main portion of the basin. Near-vertical

faults observed in the reflection images (red lines) appear to spatially correlate

with variations in the underlying basement, similar to Figure 2.7b. These faults

have been proposed to be strike-slip features associated with a prior episode

of deformation, which ceased activity ~1 Ma, as evidenced by the undisturbed

sediments overlying the faults (Tsampouraki-Kraounaki et al. 2018).

While the travel time tomographic velocity model is less detailed than the

seismic reflection images with respect to structure and deposition sequence, it

provides valuable velocity and depth constraints, as well as fuller three-dimensional

spatial coverage than the existing reflection imaging. The tomographic images also

22



0

-1

-2

-3

0 5 10 15 20 25 30

vertical exaggeration 4:1

Distance (km)

tomography - depth (km)

old margin

0.6

1.5

0 30

B B’

Updoming of Messinian Evaporites0 30

MCS - TWTT
0.6

1.5

0 30
-3.5

vertically exaggerated 4:1

Distance (km)

tomography - depth (km)

Christiana Basin

0.5

0

1

2

0 10 20 30
TW

TT
 (s

)

A A’

MCS - TWTT

0

1

2

0 10 20 30

TW
TT

 (s
)

MCS and tomography - TWTTsecondary faults MCS and tomography - TWTT

0 10 20 30

0

-1

-2

-3

El
ev

at
io

n 
(k

m
)

tomography - depth (km)vertically exaggerated 4:1

AAB

Distance (km)

AH

0.5

1

1.5

old margin

0 30

0.5

1

1.5

C C’

0 30

a) b) c)

Vp  (km/s)

Figure 2.7. Comparisons of tomography velocity model with seismic reflection
profiles. Upper panels depict seismic reflection profiles from previous studies
(locations shown in Figure 2.6). Middle panels show seismic velocity from this
study overlain on the seismic reflection image, after converting depth into a Two-
Way Travel Time (TWTT). Bottom panel is the same seismic velocity cross-section
shown as a function of depth. (a) Seismic reflection profile (H14) from Figure 6
of Nomikou et al. (2018) across the Anydros Ridge showing tomographic recovery
of the Anafi-Amorgos Basin (AAB) as well as recovery of secondary faulting on
the Anydros Horst (AH) (b) E-W seismic reflection profile across the east-central
Christiana basin (Figure 7 of Tsampouraki-Kraounaki & Sakellariou 2018) showing
the correlation between the seismic velocity structure and faults in the seismic
reflection data. The authors interpret Unit 6 as Messinian evaporites and the region
of up-doming of this layer correlates with higher seismic velocities. Red lines in
bottom panel are faults from the seismic reflection data. (c) S-N seismic reflection
profile across the eastern portion of the Christiana basin (Figure 5 of Tsampouraki-
Kraounaki & Sakellariou 2018) showing thin sediments over an old basin margin.
Shallow faults from seismic reflection studies (red lines) correlate well with deeper
undulations of basement structure.

reveal deeper features such as faults that ultimately control the shallower structures

observed in the seismic reflection images.

2.5.3 Mapping geometry and thickness of sedimentary basins from the

tomography model

We outline basin depth using a velocity of 4 km/s throughout the model

and map both the two-way travel-time (TWTT) and the depth to the 4 km/s

23



iso-velocity contour (Figure 2.8). Visual comparison between seismic reflection

images and tomography results indicates that this value typically corresponds to

the sediment-basement interface (e.g. Figure 2.7 & A.11). The 4 km/s velocity also

corresponds to the midpoint of a rapid increase in average seismic velocity with

depth (Figure 2.4), interpreted as the transition from seismically slow sediments

to the underlying, seismically fast metamorphic basement rocks. We contour the

depth to the 4 km/s iso-velocity to delineate the 3D geometry and thickness of

the sedimentary basins across the entire study area (Figure 2.8a), and not just

the profiles that were mapped using TWTT from seismic reflection studies (e.g.

Nomikou et al., 2016).

Figure 2.8 shows how the faults identified in Figure 2.5 correlate with, and

control, the geometry and internal structure of the basins. Cross-sections through

the tomography results on two SW-NE cross-sections through the Christiana Basin

west of Santorini and two WNW-ESE cross-sections through the Anydros and Anafi

Basins east of Santorini are shown in Figure 2.9 (with locations of the profiles in

Figure 2.6). This figure validates and elucidates the faults identified in the map

views (Figure 2.5) and their relationship to basin structure (Figure 2.8).

2.5.4 Western side of Santorini: Christiana Basin and Christiana Ridge

The anomalously low P-wave velocities at 0-2 km depth in the western part

of the model coincide with the NW-SE striking Christiana Basin (Figures 2.2 &

2.5). Low-velocity anomalies associated with the basin extend to ~1 km depth in

the basin center and down to 2 km in its SE and NW parts, indicating the presence

of two deeper sub-basins (Figure 2.5 and 2.9). The SE sub-basin is deepest close

to Santorini, near Christiana island (Figure 2.5d). Results from previous seismic

reflection studies suggest that these sub-basins formed in the Early to Middle

24



CB
AFB

SAB

AAB

AB

SE sub-basin  

 NW sub-basin 

thin cover of sediments

th
in

 co
ve

r o
f s

ed
im

en
ts

thin cover of 

sediments

TWTT thickness (s)

a)

b)

CB

AB

AFB

SAB

AAB

Thickness (km)

Figure 2.8. Sedimentary basin thickness maps from the seismic tomography
model. (a) Basin thickness, determined using the 4 km/s velocity contour as
the sediment-basement transition, with relevant geological features labeled. (b)
Tomographic basin thickness converted to TWTT for comparison with seismic
reflection studies. The Anydros basin, despite being a prominent tectonic feature
on bathymetric maps, has relatively shallow sediment fill (1-1.5 km). In contrast,
the Amorgos-Anafi Basin and the Santorini-Anafi basins as well as the relatively
inactive Christiana basin contain thick sediment fill (>2 km). Faults after Figure
2.5, with deeper faults (2-3 km depth) shown in light grey and shallower faults (0-1
km) in dark grey. Dark shading indicates areas with poor resolution of sediment
thickness. Labels as in Figure 2.2.

25



Christiana Basin

Christiana Basin

v (km/s)dv (km/s)

vertical exaggeration 2:1

D D’

E E’
old margin

bathymetric margin

bathymetric margin

old margin Christiana BasinCR

CR Christiana Basin

Christiana Basin
E E’

buried block

D D’
CR Christiana Basin

CR

El
ev

at
io

n 
(k

m
)

old margin

bathymetric margin

bathymetric margin

old margin

SW facing normal fault

SW facing normal fault
Distance (km)

Anydros Basin
F F’

G G’Anydros Horst

Distance (km)

Anydros Basin
F F’

G G’Anydros Horst

buried ridge

buried ridge

a) b)

Figure 2.9. Seismic velocity cross-sections through the Christiana and Anydros
basins. (a) Profiles D-D’ and E-E’ through the Christiana basin in velocity (v)
and velocity anomalies (dv) with interpretations of main fault geometries shown
in black on dv cross-sections. The open triangles mark where the profiles cross the
black faults shown on Figure 2.6. These profiles show evidence for down-dropped
blocks within the Christiana basin. (b) Profiles F-F’ and G-G’ through the Anydros
and Anafi basins, following conventions described above. These two basins show
thinner sedimentary fill than the Christiana basin, more pervasive faults at depth,
and are regionally characterized by a seismically slower basement velocities than
the profiles through the Christiana basin (D-D’ and E-E’) (see also Figure 2.4).

Pleistocene, before continued expansion was accommodated by wider faults in the

north and south (Tsampouraki-Kraounaki & Sakellariou 2018).

The Christiana basin is separated from the Sea of Crete to the south by

the Christiana ridge (Tsampouraki-Kraounaki & Sakellariou 2018) (Figures 2.2

& 2.5). The northern, bathymetrically expressed margin of the ridge proper

(dashed black line north of CR in Figures 2.5a and 2.5b) has little tomographic

expression, showing a minor velocity contrast across the fault (Figure 2.9a). The

tomographically expressed fault margin (solid black line above CR in Figure 2.5b)

is located basin-ward (north) of the ridge proper, striking NW-SE, and reflecting

the presence of an old basin margin, buried under a thin veneer of sediments

26



(Tsampouraki-Kraounaki & Sakellariou 2018, see also Figure 2.7c). This fault,

which dips at an angle of 25-40o, marks the deepening of the basin to km-scale

depths (Tsampouraki-Karounaki et al., 2018) (Figures 2.8 & 2.9a).

The northern margin of the Christiana basin is marked by a large SW-facing

normal fault. While this fault is clearly evident in Figures 2.8 & 2.9a, it has little

bathymetric expression close to Santorini (Figure 2.2), most likely because it is

covered by volcanic deposits. On the contrary, in the western part of the model

away from Santorini, it is identified in both the tomography model and the seafloor

bathymetry as seen in the D-D’ and E-E’ model cross-sections and their proposed

interpretation (Figure 2.9a). At ~2 km depth, and basin-ward of the northern

basin-bounding fault, there is a large buried normal-fault block, dipping to the

SSW at an angle of 25-40o (solid black line north of CB in Figure 2.5 and solid

black line labelled buried block in Figure 2.9a). This fault forms the northern

margin of the SE sub-basin.

2.5.5 Eastern side of Santorini: Anydros Basin, Anafi Basin and Anydros Horst

The two large NE-SW trending basins on the eastern side of Santorini,

the Anydros (AB) and Anafi (AFB) basins, are tomographically identified at 0-1

km depths as seismically slow regions (Figures 2.5a and b and Figure 2.9b). The

Pliocene/Quaternary-aged Anydros Basin has thin basin fill and appears to be only

~1.5 km thick (Figures 2.8 & 2.9b). The similarly aged Anafi basin is divided into

two sub-basins by a buried ridge at <1 km depth (Figure 2.5), the Anafi-Amorgos

basin (AAB) to the north and the Santorini-Anafi basin (SAB) to the south. Both

sub-basins are up to 2 km thick. The buried ridge dividing these sub-basins has an

ENE-WSW orientation (Figure 2.5b).

27



On the eastern side of Santorini, we observe tomographically resolved

variations in the strike of basin margins of up to 20o from a mean NE-SW

direction. For example, the northern margin of the Anydros basin, the Ios fault

(IF in Figure 2.5), strikes approximately N40oE, whereas the southern margin

of the Anydros basin strikes roughly N25oE. Similarly, the large dipping fault on

the northern margin of the Anafi-Amorgos basin (the Santorini-Amorgos fault)

is curved and has strikes ranging from N25oE to N60oE (Figure 2.5), a pattern

consistent with sea-floor morphology (e.g. Hooft et al., 2017, Nomikou et al., 2018)

(Figure 2.2b).

The Anydros and Anafi basins are separated by the Anydros Horst (AH),

which has typical bedrock velocities of 5-6 km/s at shallow depths (<1 km), in

contrast to the low velocities in the basins (2-4 km/s) (Figure 2.9b). Close to

Santorini, this horst is buried under a thin cover of sediments, possibly by volcanic

eruption deposits, as suggested by the presence of a relatively flat bathymetry but a

strong high-velocity tomographic signature.

2.5.6 Structures associated with Santorini

Several NE-SW striking faults extend from the western side of the study

area (e.g. Christiana basin region), through Santorini, to the eastern side (Anydros

Horst and Anafi basin region). At shallow depths, metamorphic blocks that are

part of the pre-existing volcanic basement and that outcrop in southern Santorini

are well recovered as seismically fast regions (Figures 2.5a and 2.10). The NE-SW

striking Akrotiri fault on the western side of Santorini, which is identified in the

bathymetry south of Santorini (Figure 2.2) (Hooft et al., 2017), is tomographically

observed at 0-1 km depth as a prominent velocity contrast that strikes through

Santorini and out to the east toward the Anydros Horst region (Figure 2.5a-b)

28



juxtaposing the seismically slow basin sediments of the Anydros Basin to the

north against the seismically fast metamorphic basement to the south. This

fault, which extends from the southwestern termination of the Christiana basin,

through Santorini, to the margin of the Anydros Horst on the east, is subparallel

to the regional alignment of volcanic centers (Figures 2.5a and 2.5b). At deeper

depths (2-3 km), a tomographically defined fault strikes through the northern

caldera basin, passing through NE Santorini. This fault may also be connected to a

bathymetrically expressed fault extending NE from Christiana island (Figure 2.2b).

From Santorini’s northern caldera, the fault extends northeast toward Kolumbo

volcano (Figures 2.5c & A.10).

At 1-2 km depth a large circular, low-velocity anomaly is observed in the

center of the northern caldera basin, which has been proposed to be related to

caldera collapse (Hooft et al., 2019) (Figure 2.10). This anomaly directly overlies

the inflation source from a 2011-2012 unrest period (inferred source depth of ~4

km) (Parks et al., 2015) and is bounded by the two dominant tectono-magmatic

lines on Santorini, the Kameni and Kolumbo lines (Pfeiffer, 2001). The southern

margin of the anomaly marks the location of recent volcanism (Kameni lavas,

Figure 2.10) (Pyle and Elliott, 2006). Its shape matches that of the shallow basin-

fill of about 100 m thick, observed in previous seismic reflection images and inferred

to correspond to LBA eruption deposits (Johnson et al., 2015).

2.6 Discussion

2.6.1 Differences between the east and west sides of Santorini and relationship to

faulting

The area east of Santorini is seismically active and exhibits ongoing,

distributed faulting including a MS = 7.4 normal-faulting event on the Santorini-

29



dv (km/s)

5 km

volcano-tectonic lines

3 km depth

approximate dike 
orientation
exposed dikes 

Kolumbo line (this study)

N

Exposed dike 
(Therasia)

local fault

Exposed dike 
orientations

N

Exposed dike 
(Therasia)

Exposed dike 
orientations

Kolumbo

Kameni Lavas

Therasia Lavas
Skaros Lavas

Peristeria Volcanics
Akrotiri Volcanics
Pre-volcanic basement

Kameni line? (this study)

1 km depth

5 km

5 km

a) b)

c)

Figure 2.10. Comparison of the seismic velocity beneath Santorini with volcanic
features and exposed geology. (a) Geologic map of Santorini (modified from Druitt
et al., 2016) with dike locations and orientations (Browning et al., 2015; Fabbro
et al., 2013), the Kolumbo and Kameni tectono-magmatic lineaments (defined by
prior studies), locations of small-scale faulting (Druitt et al., 1999), location of
the 2011-2012 inflation source (Parks et al., 2015), long wavelength gravity data
(black contours at 2 mGal intervals; Budetta et al., 1984), and a LBA (Late Bronze
Age) deposit thickness map from seismic reflection data (modified from Johnston
et al., 2015). (b) Comparison of previously proposed tectono-magmatic lineament
locations (red lines), 2011-2012 inflation point source (blue dot), and Late Bronze
Age (LBA) deposit thickness (modified from Johnston et al., 2015) with relative
velocity (dv) at 1 km depth. An increase in the LBA deposit thickness directly
overlies a decrease in velocity associated with caldera collapse (Hooft et al., 2019)
and also overlies the 2011-2012 inflation point source at ~4 km depth (Parks et
al., 2015). (c) Comparison of locations of exposed dikes, their orientations, faults,
and long wavelength gravity with dv at 3 km depth. The Kolumbo line (labelled)
is clearly observed in the tomography, dividing slower seismic velocities to the
SE from faster velocities to the NW. The orientation of the NE-SW striking line
parallels exposed dike orientations and matches the strike of regional faults. The
Kameni line (labelled) is also observed in the tomography, paralleling the Kolumbo
line.

30



Amorgos fault in 1956 (Brüstle et al., 2014; Konstantinou, 2010; Okal et al., 2009;

Papadopoulos & Pavlides, 1992). There are also two regions of small earthquakes:

those around Kolumbo that are thought to result from magmatic processes

(Bohnhoff et al., 2006; Dimitriadis et al., 2009) and clustered earthquakes around

the Anydros Horst (AH) that may reflect upward migration of fluids within a zone

of tectonic weakness or the development of volcanic activity (Bohnhoff et al., 2006)

(e.g. Figure 2.6). In contrast, the west side is characterized by the seismically

quiet, older Pliocene faulting of the Christiana Basin (e.g. Bohnhoff et al., 2006).

These differences have led to the interpretation that the AH is part of a boundary

that accommodates the relative motion between a competent Cycladic block and

a less-competent eastern Aegean region, dividing the Aegean into a seismically

quiet west and more active east (e.g., Bohnhoff et al., 2006; McClusky et al., 2000;

Reilinger et al., 2010).

At ~3 km depth the basement east of Santorini has a lower average seismic

velocity than to the west of Santorini by ~0.4 km/s (7%) (Figures 2.4 and 2.5),

which can be attributed to the presence of dense, distributed faulting in AH

region, reducing rock competency. Preliminary plots of travel time misfit as a

function of azimuth show NE-SW-oriented seismic anisotropy which results from

a fracture network in this region (Figure A.12). Focal mechanisms of shallow

earthquakes from the AH area (Friederich et al., 2014; Dimitriadis et al. 2009) are

also consistent with normal faulting due to NW-SE 130-155o extension, indicating

ongoing NE-SW fault activity in this region.

An alternative explanation for the difference in average velocity on the east

and west sides of Santorini is that it is generated by differences in composition.

While the presence of carbonates, flyschs, greenschists, blueschists, and granites

(Kilias et al., 2013) indicates that variable seismic velocities due to compositional

31



differences should exist, the orientation of the geological units is mainly controlled

by N-S-dipping regional detachments, which is unlikely to produce significant NE-

SW velocity variations. We conclude that the association of the lower P-wave

velocities east of Santorini with active deformation and pervasive fracturing of the

basement is the most plausible scenario.

2.6.2 Tectonic evolution of the basins and Christiana-Santorini-Kolumbo

magmatism: Localization of magma in a proto-Anydros Basin

Using the delineated faults and basin structures, we refine an existing model

for the tectonic evolution of the Santorini region from the Miocene-Pliocene to

present (Piper et al., 2007) and investigate how tectonism interacts with, and

localizes, magmatism (Figure 2.11). While the tomographic velocity images provide

information on the strike and style of faults, absolute time constraints on fault

activity rely on other studies, mainly seismic reflection imaging.

The earliest stage of faulting in the region, during the Miocene-Pliocene

(Figure 2.11a), resulted in WNW-ESE-oriented faults that define the Christiana

Basin and likely the buried ENE-WSW faults that divide the Anafi basin

(Anastakis and Piper, 2005; Piper et al., 2007) (hereafter these faults will be

referred to as E-W to be consistent with the literature). The offset between these

two sets of tomographically resolved faults (Figures 2.5 and 2.11a) supports the

inference of Piper et al., (2007) that a transfer zone may have existed striking

approximately N through the area where Santorini currently resides.

In the late Pliocene to Pleistocene (Figure 2.11b), these predominately

E-W structures were cross-cut by NE-SW striking faults (Piper et al., 2007).

This may have occurred as a result of NW-SE extension due to differences in

motion between the Cycladic and eastern Aegean blocks (e.g. Reilinger et al.,

32



24˚45' 25˚00' 25˚15' 25˚30' 25˚45' 26˚00'

36˚15'

36˚30'

36˚45'
24˚45' 25˚00' 25˚15' 25˚30' 25˚45' 26˚00'

36˚15'

36˚30'

36˚45'

24˚45' 25˚00' 25˚15' 25˚30' 25˚45' 26˚00'

36˚15'

36˚30'

36˚45'

24˚45' 25˚00' 25˚15' 25˚30' 25˚45' 26˚00'

36˚15'

36˚30'

36˚45'

Pliocene Pliocene Pleistocene

Pleistocene Present

Cycladic Block

Eastern Aegean Block

transfer zone (?)

a) b)

c) d)

Pro
to

-A
nyd

ro
s B

as
in

Christiana 
Volcanism

Akrotiri 
Volcanism

Kolumbo 
Volcanism

bathymetric 
margin

buried margin

buried block

Figure 2.11. Fault evolution and its relationship to magmatism: Sketch of the
evolution of faulting and volcanism in the Santorini area using the faults defined
from the tomographic model (Figure 2.5) and building on the tectonic evolution
model of Piper et al. (2007). Temporal constraints are taken from previous studies
(Piper et al., 2007). Red lines and red shading show active faults during each
time period, hatched areas indicate basins, red circles indicate regions of volcanic
activity, black arrows indicate approximate extension direction for each period. (a)
Formation of the Christiana Basin as a full-graben structure in the Pliocene under
NNE-SSW extension. (b) Transition to NW-SE extension leading to Anydros
Horst formation from the Pliocene-Pleistocene on, with initial formation of a
proto-Anydros Basin. (c) Initiation of Christiana volcanism in the Pleistocene.
(d) Volcanism at Santorini and Kolumbo developed in the Quaternary within the
proto-Anydros Basin. Current diffuse seismicity (orange region; Bohnhoff et al.,
2006, Brüstle et al, 2014) occurs predominately on the eastern side of Santorini.

2010). These faults formed primarily on the eastern side of Santorini (Figures 2.5,

2.11b) and the earlier E-W fault systems became mostly inactive (e.g. Piper et

al., 2007). The tomographic results indicate that a proto-Anydros basin formed

by progressive normal faulting along parallel to subparallel NE-SW-trending

faults near the current Anydros basin and extending further to the SW through

Santorini toward Christiana island (Figure 2.11b). A substantial portion of the

proto-basin is currently filled by the Santorini volcanic edifice (bold black line,

33



Figure 2.4c) and the Kolumbo volcanic field (Figure 2.11b, 2.11d). Gravity and

magnetic data support the presence of a NE-SW-oriented fault zone underlying

the caldera (Figure 2.10) (Budetta et al., 1984). We infer that the proto-Anydros

basin intersected the Christiana basin in the region of the SE Christiana sub-basin,

potentially contributing to deepening of the SE portion of this sub-basin (Figure

2.8).

During the Pleistocene (Figure 2.11c), volcanism initiated at Christiana

Island (~1.2 Myr) (Piper et al., 2007) and was roughly localized at the intersection

of the proto-Anydros basin with the Christiana basin (Figure 2.11c). Because the

proto-Anydro basin is filled by the Santorini and Kolumbo edifices (Figure 2.11d),

the NW-SE extension that formed this basin must have preceded the majority

of these volcanic eruptions. Volcanism at Santorini began around the Middle

Pleistocene (650 ka) (Figure 2.11c) (Druitt et al., 1999). Seismic reflection data

show that the present-day Anydros basin opened in six tectonic pulses (Hübscher et

al., 2015; Nomikou et al., 2016), creating the basin that is currently observed to the

NE of Santorini. The exact relationship between Santorini volcanism and Anydros

basin formation is not known though the seismic reflection data indicate that

eruptions from nearby Kolumbo seamount started after Anydros basin initiation

(Nomikou et al., 2018). Because all of the regional volcanism falls within the

proto-Anydros basin (Figure 2.11d), it is likely that this basin has played a role

in localizing volcanism. Such interaction between extension and volcanism is readily

observed in analog models where volcanism is localized in rift basins during oblique

extension (Corti et al., 2003).

The above interpretations, where proto-Anydros basin formation precedes

major local volcanism, are consistent with previous models that inferred that the

rotation from older NNE-SSW extension to the current NW-SE extension in the

34



early Pliocene was associated with an increase in volcanism in the area (Piper et

al., 2007). Across the Hellenic arc, volcanism is dominantly associated with NE-

SW striking features (i.e. NW-SE extension) (e.g., Budetta et al., 1984) and only

secondarily with NW-SE strikes (Kokkalas & Aydin, 2013).

2.6.3 Volcano-tectonic lineaments

At Santorini we recover and improve on the geometry of the well-known

Kameni and Kolumbo volcano-tectonic lineaments, using information inferred

from the seismic tomography images (Figures 2.5, 2.8 and 2.10). The northern

margin of the low-velocity anomaly (LVA ) within the Santorini caldera is bordered

by the Kolumbo line, initially identified as an alignment of volcanic edifices in

the northern portion of Santorini. It has been suggested that one of the vents of

the LBA eruption was located along this line (Pfeiffer, 2001). In our study, the

Kolumbo line is tomographically imaged at 3 km depth as a linear feature that

divides a seismically fast region to the north from a lower velocity region to the

south, oriented more northerly than suggested by previous studies and sub-parallel

to the regional NE-SW trending faults (Figures 2.2 and 2.10). The tomographically

defined Kolumbo lineament is associated with faulting observed in MCS profiles

north of Kolumbo volcano (Figure A.10) (Nomikou et al., 2016). This line is also

aligned with exposed dikes in the northern portion of the caldera walls and with

a similarly oriented dike exposed on Therasia on the western side of the caldera

(Browning et al., 2015; Fabbro et al., 2013). The joint interpretation of these

observations suggests that (at depth) the Kolumbo line is controlled by regional

tectonism and that it plays a dominant role in localizing magmatic and volcanic

features, both within, and close to, Santorini.

35



In contrast to the Kolumbo line, the Kameni line has a weaker tomographic

signature. This line, which is thought to have focused vents of prior explosive

eruptions (Druitt et al., 1989), has previously been defined using the orientation of

post-LBA eruption vents on the Kameni islands (Pyle and Elliott, 2006). Moreover,

it was re-activated seismically in the recent 2011-2012 seismo-volcanic crisis (e.g.

Newman et al., 2012), verifying its active character. Our study shows that the line

is associated with a NE-SW trending region of limited lateral extent that divides

anomalously slow velocities to the NW from faster velocities to the SE (Figure

2.10) and lies parallel to the tomographically recovered Kolumbo line.

The Kameni and Kolumbo lines border the caldera LVA, a region of collapse

in the northern caldera (Hooft et al., 2019). This isolated region of deepened

caldera collapse, which overlies a recent influx of magma (e.g. Newman et al.,

2012; Parks et al., 2015), seems to have been limited in extent by these tectono-

magmatic lineaments. Similarly, the majority of exposed dikes and vents associated

with Santorini volcanism also fall between the Kolumbo and Kameni lines. It is

clear these lineaments have helped to control the tectonic evolution of the volcanic

system and, given their association with eruptive vents, they have likely shaped

magma input into the upper crust.

2.7 Conclusions

We used the seismic velocity model to outline faults and basins (Figure 2.5)

and to correlate magmatic and tectonic processes. In the Santorini region, we find

that the tomographic results compare well with previous seismic reflection studies

and geological outcroppings on Santorini (Figure 2.10). Geologically observed dikes

and volcanic chains around the volcano directly overlie tomographically observed

faults and tectono-magmatic lineations. Inside the caldera, volcanic vents and

36



deposits from the Late Bronze Age eruption correspond to a seismically slow

region of caldera collapse. Our results and interpretations, combined with those

from other studies, support the following conclusions: 1. Resolved tomographic

structures, specifically orientations of faults and basins, are consistent with a

previously proposed tectonic evolution model consisting of NNE-SSW extension

during the Miocene-Pliocene transitioning to NW-SE extension in the Late

Pliocene-Pleistocene that predominates on the eastern side of Santorini (Figure

2.11). This present-day extension is associated with pervasive faulting east of

Santorini. 2. The Christiana, Santorini, and Kolumbo volcanic centers are found

to lie in a NE-SW-trending basin-like structure, the proto-Anydros basin (Figure

2.5). These results are in agreement with models that suggest volcanism initiated

after the transition to NW-SE extension in the Pleistocene. 3. Two tectono-

magmatic lineaments control magma emplacement at Santorini. These lineaments

border a NE-SW trending region of low velocity, are associated with dikes in the

caldera walls, and strike parallel to active NE-SW trending regional faults and

sub-parallel to regional volcanic chains. The two lineaments also bound a region

of localized caldera collapse within Santorini’s northern caldera, verifying the strong

relationship between tectonics and magmatism.

2.8 Bridge

In Chapter II I outlined regional tectonic and magmatic features around

Santorini Volcano, emphasizing that the magmatic features, which generally have

a NE-SW orientation, have been localized in a buried NE-SW elongate basin. I

proposed a simplified conceptual model for regional tectonic evolution and its

association with this localized magmatism. In Chapter III, I discuss whether this

localized magmatism results from small scale (km) faulting or regional tectonic

37



stresses. In addition, I investigate pervasive regional fracturing of the crust in this

extensional tectonic environment.

38



CHAPTER III

RELATIONSHIP BETWEEN FAULTING AND MAGMATISM AROUND

SANTORINI: SEISMIC ANISOTROPY RESULTS FROM A REFRACTION

EXPERIMENT

This chapter is being prepared for future publication. Ben Heath conducted

all inversions, drafted all figures (excluding portions of those credited to other

sources) and wrote the manuscript. Emilie Hooft provided edits to a previous

version of this chapter. Michele Paulatto and Joanna Morgan aided anisotropy

interpretations as well as provided suggestions on the modeling of anisotropy.

3.1 Introduction

Investigations of magmatic centers and their evolution show strong evidence

that extensional tectonic and magmatic processes interact on regional ( ≥ O
(
101

)
km) and local (< O

(
101

)
km) scales. For example, volcanoes are often found to be

located in areas of increased faulting near first-order, regional fault zones. In rifts,

magmatism is often focused in regional-scale basins and vents are arranged in linear

lines paralleling the orientation of fault zones (e.g. Casey et al., 2006; Muirhead et

al., 2015). For areas such as the Basin and Range and the Aegean, magmatism

is often found to focus at accommodation zones, complex zones of intermeshed

faults (Faulds & Varga, 1998). In such regions, major normal faults often terminate

in magmatic centers, either due to pre-existing magmatism inhibiting fault

propagation or due to tectonic-induced stress distributions favorable to magma

body development (Faulds & Varga, 1998). These observations all support a strong

correlation between extension in faulted continental crust and magmatic activity.

Despite this strong correlation, understanding of exactly how these processes

are related remains enigmatic. For example, faults and fractures are thought

39



to represent permeability pathways through the crust and/or represent zones

of crustal weakness characterized regionally by damaged host rock. In these

models, dikes and their associated melt follow preexisting fracture networks, and

therefore are focused to regions dominated by extensive fracturing, extensive

damage and higher permeability (Kokkalas & Aydin, 2013). Other models suggest

that magmatism accommodates extensional strain, inhibiting faulting (Parsons

& Thompson, 1991) and that magmatism should therefore be associated with

lessened faulting. Stress-based models suggest that dikes open approximately in the

direction of least compressive stress, a direction that need not be associated with

the shear-based faulting (Anderson, 1951; Gudmundsson, 1995). In these models,

dikes follow solely a stress-based path and, therefore, a regional correlation between

faulting and magmatism represents the shared dependence of both processes on

stress. These processes are further complicated by the fact that magmatism and

tectonism iteratively impact each other. For example, dikes can alter a tectonic

stress field, leading to shallower angle faulting (Parsons & Thompson, 1993),

which in turn alters the stress field leading to altered dike orientations. From

observations it is clear that tectonism and magmatism are related, however why

they are related is still an area of active research with models proposing multiple

styles of interaction between tectonic and magmatic features.

Here we use anisotropic P-wave travel-time tomography to investigate local

and regional interaction between tectonism and magmatism at Santorini Volcano

Greece. We investigate tectonic and magmatic sources of anisotropy, modeling

the role of each in our observations. We highlight the relationship between local

faults/fractures, regional tectonic features and magmatic features. Our results show

how pervasive regional faulting has shaped the Aegean crust around Santorini.

These results support the hypothesis that regional scale tectonic processes

40



dominate the relationship between tectonism and magmatism in extensional

volcanic arcs such as the Aegean and that local-scale surface faulting neither

controls, nor is controlled by, regional magmatic activity and is rather a product

of local stress variations in the upper crust.

3.2 Background

3.2.1 Geologic Background of the Aegean

The Aegean is characterized by extension dominated tectonic structures

resulting from Oligocene/Miocene to present slab roll-back of the Hellenic

subduction zone (Pichon & Angelier, 1979), westward extrusion of Anatolia

(McKenzie, 1972), and gravitational collapse of continental crust (McKenzie, 1972).

These structures weakened the crust, localizing deformation (Jolivet et al., 2013,

and citations therein) and magmatism in the Cyclades (Kokkalas & Aydin, 2013)

(Figure 3.1).

Extensional detachments and NNE-trending strike-slip faults controlled

Aegean pluton emplacement during the middle Miocene (Kokkalas & Aydin,

2013). In many cases, this pluton emplacement was associated with structural

complexities that enhanced local permeability. This regional pluton emplacement,

which occurred along zones of contemporaneous deformation, arrested in the late

Pliocene when internal deformation of the Cycladic region ceased. Santorini lies at

the margin of this non-deforming region (e.g. Le Pichon & Kreemer, 2010).

Around Santorini during the Miocene to Pliocene, N-S extension created E-

W-striking normal faults (Anastasakis & Piper, 2005; Piper et al., 2007). Activity

on the main bounding faults of the basin ceased 1 Ma during the Pleistocene

(Piper & Perissoratis, 2003) but some regional tectonic activity continues to present

day as evidenced by small-offset faults (Tsampouraki-Kraounaki & Sakellariou,

41



25˚00' 25˚15' 25˚30' 25˚45' 26˚00'

36˚15'

36˚30'

36˚45'

−1200 −800 −400 0 400

20˚ 22˚ 24˚ 26˚ 28˚ 30˚
32˚

34˚

36˚

38˚

40˚

42˚

Methana

Milos
Santorini

Nisyros&Kos

Aegean

Greece

Turkey

Africa

Hellenic Trough

North Anatolia Fault

Cyclades

Anatolia

Crete

Cretan Basin

Eastern Aegean

volcanic arc

volcano

study area

region of 
volcanism

stretching 
lineation

burie
d horst

acc
ommodatio

n zo
ne

Figure 3.1. a) Regional map of the Aegean showing the Hellenic subduction
zone, Hellenic volcanic arc, as well as the Santorini study area, from Heath et al.
(2019). Bold black lines represent regional faults. b) Topographic/bathymetric
map showing the experiment geometry, shot locations (red dots) and seismic
stations(yellow). Orange shaded regions mark regions of magmatism. White shaded
arrows and lines represent stretching lineations recorded in metamorphic rocks
from prior episodes of deformation (after Schneider et al. (2018). Red lines mark
tectono-magmatic lineaments.

42



2018). These E-W striking normal faults are to first