University of Florida Thesis Or Dissertation

INVESTIGATION OF THE MAULE, CHILE RUPTURE ZONE USING SEISMIC ATTENUATION TOMOGRAPHY AND SHEAR WAVE SPLITTING METHODS

MEGAN ELIZABETH TORPEY

A DISSERTATION PRESENTED TO THE GRADUATE SCHOOL OF THE UNIVERSITY OF FLORIDA IN PARTIAL FULFILLMENT OF THE REQUIREMENTS FOR THE DEGREE OF DOCTOR OF PHILOSOPHY

UNIVERSITY OF FLORIDA

2016

To my family, for always being there for me and supporting me throughout this journey

ACKNOWLEDGMENTS

I would first like to express my appreciation to my advisor, Raymond Russo, who provided me with amazing opportunities throughout my graduate career. I am also thankful for his guidance and motivation in my research projects, without which it would not have been possible for me to succeed. I would like to thank my committee members

Mark Panning, David Foster, Liz Screaton, James Channell, Tim Olson, and Jean

Larson, who all helped me to better understand my research projects. Additionally, I would like to thank Sebastien Chevrot for hosting me at the Observatoire Midi-Pyrenees in Toulouse, France and guiding my research as part of the Chateaubriand Fellowship. I would also like to thank my fellow “seismo” group members, Christian Stanciu, Paul

Bremner, Sutatcha Hongsresawat, Stephanie James, Matt Farrell, and Emily Rodriguez who were always willing to lend a helping hand, whether it was during fieldwork or in the office. I would like to especially thank the office staff of the UF Department of Geological

Sciences, Pamela Haines, Nita Fahm, and Carrie Williams who made it possible for me to remain on track with academic paperwork. And lastly, I would like to thank the

National Science Foundation for funding this grant (EAR-1045609).

TABLE OF CONTENTS

page

ACKNOWLEDGMENTS ...... 4

LIST OF TABLES ...... 7

LIST OF FIGURES ...... 8

LIST OF ABBREVIATIONS ...... 11

ABSTRACT ...... 12

CHAPTER

1 INTRODUCTION ...... 14

1.1 Motivation for the Project ...... 14 1.2 2010 Mw 8.8 Maule, Chile Earthquake ...... 16 1.3 Tectonic Setting ...... 18 1.3.1 Tectonic Overview ...... 18 1.3.2 Convergence Rate and Nazca Age ...... 20 1.3.3 Subduction Dip ...... 20 1.3.4 Bathymetric Features ...... 22 1.3.5 Crustal Shortening ...... 23 1.3.6 Maule Rupture Zone ...... 23

2 SEISMIC ATTENUATION WITHIN THE MAULE, CHILE RUPTURE ZONE ...... 32

2.1 Introduction ...... 32 2.2 Seismic Attenuation Background ...... 36 2.3 Data and Methods...... 38 2.4 Results ...... 44 2.5 Model Resolution ...... 45 2.6 Discussion ...... 46 2.6.1 Large-Scale Subduction Structures ...... 46 2.6.2 High-Qs Bodies at the Surface ...... 47 2.6.3 Cobquecura Anomaly ...... 48 2.7 Conclusion ...... 51

3 SHEAR WAVE SPLITTING OF TELESEISMIC EVENTS ...... 81

3.1 Background ...... 81 3.2 Previous Teleseismic Shear Wave Splitting Studies in South America...... 81 3.3 Data and Methods...... 84 3.4 Results ...... 88 3.5 Discussion ...... 88

3.5.1 Time Delay ...... 88 3.5.2 Contributions from Crustal Anisotropy ...... 89 3.5.3 Fossilized Lithospheric Fabrics ...... 89 3.5.4 Trench Parallel Splitting and Sub-Slab Mantle Flow ...... 90 3.5.5 Alignment with Nazca Slab Depth Contours ...... 91 3.6 Conclusion ...... 92

4 SHEAR WAVE SPLITTING OF MAULE AFTERSHOCKS ...... 104

4.1 Background ...... 104 4.2 Data and Methods...... 104 4.3 Results ...... 106 4.4 Discussion ...... 106 4.4.1 Supra-slab Entrained Mantle Flow ...... 107 4.4.2 Contributions from Crustal Anisotropy ...... 107 4.4.3 Lateral Mantle Flow in the Forearc ...... 110 4.4.4 Robustness of Multichannel Analysis Method for Local Datasets ...... 110 4.4.5 Comparison to Teleseismic Dataset ...... 111 4.6 Conclusion ...... 112

5 SUMMARY AND CONCLUSIONS ...... 122

APPENDIX

A SEISMIC ATTENUATION THEORY ...... 127

A.1 Sources of Attenuation ...... 127 A.2 Quantifying Attenuation ...... 127

B ANISOTROPY AND SHEAR WAVE SPLITTING THEORY ...... 130

LIST OF REFERENCES ...... 134

BIOGRAPHICAL SKETCH ...... 151

LIST OF TABLES

Table page

3-1 Teleseismic event locations ...... 94

3-2 Teleseismic splitting measurements ...... 95

4-1 Local event locations ...... 113

4-2 Local splitting measurements ...... 116

LIST OF FIGURES

Figure page

1-1 Historic earthquakes from Hayes et al. (2013)...... 26

1-2 IRIS CHAMP temporary broadband seismic network...... 27

1-3 CHAMP stations with alphanumeric labels ...... 28

1-4 From Kroner and Stern (2005), map of Gondwana at the end of the Neoproterozoic (~540 Ma) showing the arrangement of Pan-African belts...... 29

1-5 Modified from Ramos (1988), terranes accreted to Southern South America and South American cratonic regions...... 30

1-6 Modified from Hicks et al. (2014), the large-scale geologic units of the Maule rupture zone (33°S-38°S)...... 31

2-1 Main morphostructural units in study area...... 53

2-2 The P, S, and pre-signal noise windows are shown to demonstrate the evolving time window method for the April 5, 2010 seismic recording ...... 54

2-3 Illustration of the evolving time window method...... 55

2-4 Frequency spectra for event shown in Figure 2-2 and Figure 2-3 ...... 56

2-5 Illustration of the ray path distribution in the Maule rupture volume. Colors indicate different Q values for the respective ray path...... 57

2-6 Illustration of the tomography concept...... 58

2-7 Plot of data misfit versus model residuals for different damping values tested in the inversion process...... 59

2-8 Study area for the Maule attenuation tomography ...... 60

2-9 Cross section A-A’ through attenuation tomography model...... 61

2-10 Cross section B-B’ through attenuation tomography model...... 61

2-11 Cross section C-C’ through attenuation tomography model ...... 62

2-12 Cross section D-D’ through attenuation tomography model...... 62

2-13 Cross section E-E’ through attenuation tomography model ...... 63

2-14 Cross section F-F’ through attenuation tomography model ...... 63

2-15 Depth slice at 5 km...... 64

2-16 Depth slice at 14 km...... 65

2-17 Depth slice at 23 km...... 66

2-18 Depth slice at 33 km...... 67

2-19 Depth slice at 42 km...... 68

2-20 Depth slice at 51 km...... 69

2-21 Ray path hit count for depth 0 – 9.3 km...... 70

2-22 Ray path hit count for depth 9.3 – 18.7 km...... 71

2-23 Ray path hit count for depth 18.7 – 28.0 km ...... 72

2-24 Ray path hit count for depth 28.0 – 37.3 km...... 73

2-25 Ray path hit count for depth 37.3 – 46.7 km...... 74

2-26 Input Q model (checkered Q=200 and Q=500) and output model after resolution testing...... 75

2-27 Input Q model (checkered Q=200 and Q=500) and output model after resolution testing at -34.1°S, -35.1°S, and -36.2°S latitude slices...... 76

2-28 Cross sections used in Hicks et al. (2014) velocity model...... 77

2-29 Comparison of attenuation tomography model to velocity tomography model. .. 78

2-30 Comparison of attenuation tomography model to velocity tomography model and gravity...... 79

2-31 Attenuation model at 23 km depth with high frequency rupture propagation dots from Kiser and Ishii (2011) superimposed...... 80

3-1 Schematic representation of Russo and Silver (1994) of retrograde flow model for the Nazca-South America subduction zone...... 97

3-2 Teleseismic event locations...... 98

3-3 Radial and transverse component records to illustrate Chevrot (2000) shear wave splitting intensity method ...... 99

3-4 Zooming in on the window selected in Figure 3-3...... 99

3-5 Backazimuth vs SI plot...... 100

3-6 Rose histogram of the teleseismic fast directions for the entire dataset...... 100

3-7 Teleseismic shear wave splitting measurements from SK(K)S phases recorded at stations in the IRIS CHAMP temporary seismic network ...... 101

3-8 Histogram of splitting directions ...... 102

3-9 Shear wave splitting measurements plotted with depth contours to the top of the subducting Nazca lithosphere ...... 103

4-1 Local splitting measurements from upgoing S waves recoded at IRIS Champ temporary seismic network, denoted by black triangles...... 117

4-2 Rose histogram of fast directions from the entire local dataset...... 118

4-3 Local fast polarization directions colored by direction...... 119

4-4 Schematic diagram of model from Long and Silver (2008)...... 120

4-5 Local S splitting measurements from MacDougall et al. (2012)...... 121

B-1 Incident shear wave pulse travels through anisotropic layer and is split yielding two transmitted pulses ...... 133

B-2 Schematic of an SKS and SKKS ray path ...... 133

LIST OF ABBREVIATIONS

APM Absolute Plate Motion

BANJO Broadband Andean Joint Experiment

CA Cobquecura Anomaly

CHAMP Chile Rapid Aftershock Mobilization Program

CHARGE Chile Argentina Geophysical Experiment

IMAD International Maule Aftershock Deployment

IRIS Incorporated Research Institutions for Seismology

Moho Mohorovicic discontinuity

PISCO Projecto de Investigacion Sismologica de la Cordillera Occidental

Qs Shear wave attenuation quality factor

SEDA Seismic Exploration of the Deep Altiplano

SI Splitting Intensity

TIPTEQ The Incoming Plate to Mega-Thrust Earthquake Processes

USGS United States Geological Survey

Abstract of Dissertation Presented to the Graduate School of the University of Florida in Partial Fulfillment of the Requirements for the Degree of Doctor of Philosophy

INVESTIGATION OF THE MAULE, CHILE RUPTURE ZONE USING SEISMIC ATTENUATION TOMOGRAPHY AND SHEAR WAVE SPLITTING METHODS

Megan Elizabeth Torpey

May 2016

Chair: R.M. Russo Major: Geology

The Maule, Chile 2010 Mw 8.8 earthquake afforded the opportunity to study the rupture zone (33°S-38°S) in detail using aftershocks recorded by the rapid-response

IRIS CHAMP seismic network. We used measurements of differential S to P seismic attenuation to characterize the attenuation structure of the South American crust and

-1 upper mantle wedge. We implemented an evolving time window to determine Qs values using a spectral ratio method and incorporated these measurements into a

-1 bounded linear inequality least squares inversion to solve for Qs in a 3D volume. On a large-scale, we observe an east-dipping low attenuation feature, consistent with the location of the Nazca oceanic slab, and image progressively greater attenuation as we move towards the surface of our model. A dramatic feature in our model is a large, low- attenuation body in the same location where Hicks et al. (2014) resolved a high P wave velocity anomaly in their velocity tomography model.

We calculated the shear wave splitting intensity of the Maule rupture zone by implementing the multichannel method of Chevrot (2000) which calculates the splitting intensity of teleseismic SK(K)S phases and splitting parameters,  and t. The results we obtained show an overall fast direction with a strong component of trench parallel

splitting and very few trench normal splits. The fast directions do not parallel the Nazca

APM, but are instead dominated by splits rotated 40°-50° counter-clockwise from Nazca

APM. Based on these data, we see little evidence for sub-slab entrained mantle flow and invoke the trench-parallel retrograde flow model as an explanation for our measurements.

We developed an extended splitting intensity method to allow for use of the upgoing S phase from Maule aftershocks, utilizing the initial event polarization. For this local dataset, we observe three dominant fast directions oriented N20°W, N40°E, and

N10°W-20°E and a subset of fast directions trending N60°-90°E which suggests that the local splits are generated in rocks strongly deformed by shear due to plate convergence, but also may be a result of flattening through conjugate shearing in crustal rocks. Additionally, splits oriented N10°W-20°E are nearly strike-parallel and could indicate supra-slab trench-parallel flow.

CHAPTER 1 INTRODUCTION

1.1 Motivation for the Project

The South American subduction zone is a very seismically active margin and has been host to many devastating earthquakes (e.g. Beck et al., 1998). Rupture segmentation along this margin, however, has been very heterogeneous and can be largely attributed to heterogeneities present on the subducting Nazca oceanic lithosphere and in the overriding South American continental crust. Bathymetic highs and lows on the subducting Nazca plate, such as seamounts and ridges, introduce heterogeneities at the subduction trench that can as either barriers to rupture, which hinder the propagation of seismic energy, or as asperities, which may augment or even induce seismogenesis. Additionally, changes in the Nazca slab morphology along-strike have a large effect on the overall working of the subduction zone. Two flat-slab subduction segments along the subduction zone, the Peruvian segment and the

Argentina-Chile segment, dramatically affect the deformation in the overriding South

American plate. Above both the flat-slab segments, the South American lithosphere is less deformed and lacks active volcanism, whereas in regions of normal-slab subduction, there is large-scale deformation and volcanism present. Heterogeneities such as these can act to segment rupture and, to better understand seismogenesis in subduction zones, it is important that we do high-resolution studies of subduction systems when possible to assess the presence of potential rupture barriers or asperities.

In February of 2010, a Mw 8.8 megathrust earthquake occurred along the South

American margin near the Chilean town of Maule, partially filling the ‘Darwin seismic

gap’, a segment of the subduction zone that had not slipped co-seismically in its entirety since the Mw 8.5 earthquake in 1835 (Hayes et al., 2013). The Maule earthquake is the sixth largest event recorded on modern seismological instrumentation. As a response to this earthquake, an international team of scientists deployed a rapid-response temporary seismic network, the International Maule Aftershock Deployment (IMAD), to capture the thousands of anticipated subsequent aftershocks. The motivation for this study came from the copious dataset of aftershocks recorded by IMAD, which presented an excellent opportunity to study the Maule rupture zone in great detail. In this dissertation, I present three studies aimed at investigating the rupture zone (33°S-

38°S) of the 2010 Maule, Chile earthquake. In Chapter 2, I investigate the seismic shear wave attenuation structure of the subduction system in the Maule segment, because shear waves are very sensitive indicators of rheology and will allow us to assess the rheological boundaries that may act to segment rupture as barriers or asperities. In

Chapters 3 and 4, I determine the shear wave splitting intensity recorded within the rupture zone from teleseismic events (Chapter 3) and local events (Chapter 4). Shear wave splitting reveals information about anisotropy in the medium and can be used as a proxy to learn about tectonic processes. Because of the different wavelengths, teleseismic shear wave splitting is sensitive to large-scale processes whereas local shear wave splitting is much more sensitive to deformation on a much smaller scale.

These shear wave splitting measurements can provide an ancillary dataset for our attenuation model. All data used in this study were recorded at stations of the IRIS

CHAMP seismic network, a subset of the IMAD stations that included uniform high- quality broadband instrumentation recording at 100 Hz sampling.

1.2 2010 Mw 8.8 Maule, Chile Earthquake

On February 27th of 2010, a megathrust earthquake occurred along a portion of the Nazca-South American subduction zone, northwest of metropolitan Concepción in central Chile, with an epicenter, as determined by the International Seismological

Centre (ISC), at 36.17°S and 72.84°W and at a depth of ~35 km (Lew et al., 2010). The moment magnitude (Mw) of the Maule earthquake main shock was determined by the

USGS to be Mw 8.8. The earthquake rupture length along the Nazca-South America plate boundary was approximately 450-550 km (Lay et al., 2010), extending from just south of Valparaiso to the Arauco Peninsula. In addition to this long rupture length parallel to the coast, the USGS reported a 100 km rupture width. The best double- couple fault plane solution for the event strikes N18°E, dips 18°E, and the slip vector rake is 112°, indicating nearly pure thrusting of South America over the Nazca plate

(Lay et al., 2010). The USGS centroid moment tensor (CMT) solution is consistent with large historic earthquakes along the Chilean subduction zone (Bilek 2009), and indicates that the earthquake occurred on a shallow, low-angle thrust fault, i.e. the shallow portion of the Nazca-South America interplate interface (Delouis et al., 2010).

Focal mechanisms of larger magnitude aftershock seismicity show a predominance of thrust faulting, with the majority of these events located 40 to 140 km east of the trench. Ryder et al. (2012) determined aftershock depths to be consistent with slab depths of the global slab model SLAB1.0 (Hayes et al., 2012) in the northern and southern portions of the rupture zone, but found seismicity in the central portion of the rupture zone to be concentrated above the nominal interplate interface of the

SLAB1.0 model.

Kiser and Ishii (2011) back-projected the high frequency seismic energy recorded at stations in the United States and Japan to determine the rupture propagation of the

Maule earthquake. They found that high frequency seismic energy that defines the rupture front radiated from SW to NE along the interplate interface, with a rupture velocity increasing with distance from the hypocenter from 2.2 km/s around 36.5°S latitude to 2.9 km/s near 34.5°S latitude. A minor sequence of rupture propagation to the

SE of the hypocenter occurred at a much lower velocity of 0.8 km/s near 38°S. A second back-projection study by Koper et al. (2012) found results largely consistent with those of Kiser and Ishii (2011), resolving somewhat different regions of overall peak energy release at all frequencies.

A number of co-seismic slip inversion studies, as well as joint inversion studies incorporating seismic, InSAR, spaceborne gravimetry, geodetic, and/or tsunami data, consistently show two zones of greater slip, one to the north of the epicenter (34°S-

36°S) and a smaller zone to the south (37°S-38°S) (Tong et al., 2010; Lay et al., 2010;

Delouis et al., 2010; Lorito et al., 2011; Vigny et al., 2011; Pollitz et al., 2011; Moreno et al., 2012; Wang et al., 2012; Lin et al., 2013). Although most slip models entail a consistent extent of slip in the northern region, many vary in their assessment of both the extent of slip and the depth of maximum slip in the southern region. The analyses of

Moreno et al. (2010, 2012) of the pre-seismic locking distribution in the rupture zone show two primary regions of high coupling that coincide with the two aforementioned regions of slip to the north and south. Lorito et al. (2011), however, asserted that a high degree of pre-seismic locking is not sufficient to explain the large slip of the Maule

event, and that other factors must be considered, such as fault rheology, structural heterogeneity, and aseismic slip.

Gravimetric studies following the megathrust event (e.g. Han et al., 2010; Heki and Matsuo, 2010) used the GRACE satellite system to observe regional gravity changes following the 2010 Chile earthquake. These studies document a decrease of 5

Gal east of the Maule epicenter with a wavelength of 500 km (Han et al. 2010). Han et al. (2010) attribute this change in mass distribution to be due to both crustal dilatation and surface subsidence following the megathrust.

1.3 Tectonic Setting

1.3.1 Tectonic Overview

The accretion and amalgamation of various terranes of Archaean-Proterozoic age during the Pan African-Brasiliano Orogeny was responsible for the assembly of the super continent Gondwana, which included the South American block, composed of three main cratons: Amazonia, Rio de la Plata and Sao Francisco (Ramos, 1988; Brito

Neves et al., 1999). During the early Paleozoic, diverse allochthonous terranes accreted to the proto-margin of Gondwana as a result of eastward subduction of the oceanic crust of the ancestral Pacific (Iapetus) oceanic plate, microplate collision, and terrane accretion (Alvarez et al., 2011). One of the terranes that accreted to the western margin of Gondwana was the Chilenia Terrane (Ramos et al., 1984, 1986); Chilenia rocks now form the crustal basement of South America in the Maule rupture region. Tectonic reconstructions suggest that the Chilenia Terrane accreted to the proto-margin of

Gondwana in the Late Devonian to Early Carboniferous, i.e., ~420-390 Ma (Ramos

1994; Ramos 1996; Keppie and Ramos 1999). In the Late Devonian, the Patagonia

Terrane accreted to Gondwana and constitutes the southern region of present day Chile

(Keppie and Ramos 1999). Polar migration curves for South America and Africa (Vilas and Valencio 1978) suggest that there was a period following the Late Paleozoic in which Gondwana’s motion relative to the Earth’s magnetic pole largely ceased. Possible causes for this deceleration include the closure of the Iapetus ocean and subsequent collision of Gondwana with Laurentia (Charrier et al., 2007).

During this stationary period, continued subduction beneath Chilenia resulted in heat accumulating in the upper mantle beneath at least this portion of Gondwana’s western margin, inducing crustal softening and magma production, generating the

Coastal Cordillera as a composite volcanic arc and batholith complex constructed on and intruded into Chilenia (Charrier et al., 2007). Following the disaggregation of

Gondwana in the Late Paleozoic, the western South American margin evolved into an ocean-continent subduction zone characterized by voluminous felsic magmatism and related extensional rifting (Ramos and Aleman, 2000). An accretionary prism developed along the Chilenia margin during the Late Carboniferous to Late Triassic, and rocks of both shallow, lightly metamorphosed portions of this accretionary complex, and of a more highly metamorphosed, more deeply subducted accretionary rocks are currently exposed in the Coastal Cordillera. Tectonic erosion removed significant volumes of the

Late Paleozoic subduction complex along Chilenia’s western margin (Mpodozis and

Ramos 1990; Willner et al., 2005). A rift system developed in the Triassic, causing extension with a general NW-SE trend, heavily controlled by previously developed basement fabrics. This rifting was accompanied by production of widespread mixed composition magmatism, including anorogenic (a-type) granitoids, including olivine, derived from melting of accretionary prism rocks in close association with melting

mantle rocks (Vásquez et al., 2011). These Triassic intrusives are now exposed along the Coastal Cordillera, and become more common at the surface from Cobquecura northwards.

During the Cretaceous ~134 My, oceanic magnetic anomalies record the beginning of a drift phase as South America separated from Africa during the opening of the southernmost portion of the Atlantic Ocean (Mueller et al., 1998), with propagation northward toward the Equatorial Atlantic Ocean (Channell et al., 1995).

1.3.2 Convergence Rate and Nazca Age

As a result of the separation of the Farallon plate into the northern Cocos and southern Nazca plates, around 32-25 Ma, a period of higher and more orthogonal convergence began along the Chilean subduction zone (Pardo-Casas and Molnar,

1987). Currently, at the latitudes of the Maule rupture zone, the convergence azimuth between the Nazca plate and South American plate is ~78°E, with a geologically- derived convergence rate of ~74 mm/yr (DeMets et al., 2010) and a slightly slower rate deriving from GPS observations (Altamimi et al., 2007; Vigny et al., 2010). The Nazca plate decreases in age from 45 Ma just north of the Maule rupture zone to 0 Ma at

45.5°S, where the Chile Ridge subducts (Cande and Leslie 1986; Russo and Silver,

1996; Russo et al., 2010a,b); the age of the Nazca plate that underthrusts South

America during the Maule main shock decreases from ~37 Ma to 28 Ma north to south along the rupture zone (Mueller et al., 1998).

1.3.3 Subduction Dip

In addition to the variation in age of the Nazca lithosphere subducting along the

South American margin, the dip of the slab changes abruptly several times along strike

(Cahil and Isacks 1992), forming folds with down-dip fold axes of the entire subducting

lithosphere, resulting in an apparent segmentation of the subduction zone. Beneath the

Northern Andes (8°S-14°S) in Peru, the Nazca plate subducts at a shallow angle and then flattens beneath a significant portion of the overriding plate (and high Andes), before resubducting to the bottom of the upper mantle transition zone (Cahill and

Isacks, 1992), where its presence is inferred from steady, deep seismic moment release. The Peru flat slab region is marked by an absence of arc volcanism, in strong contrast to the voluminous extrusive magmatism associated with the subduction zone to the north and south of the flat slab region (Kay et al., 1991; Ramos et al., 1991; Kay and

Abbruzzi, 1996). In the Arica bend region of the Central Andes (14°S-27°S), the Nazca slab subducts at an angle of ~30°, and a vigorously active continental arc lies above this portion of the subduction zone. To the south (27°S-33°S), including the northern portion of the Maule rupture zone, the Nazca slab is again folded, forming a second flat slab segment (Anderson et al., 2007). Southward of this flat slab (33°S-40°S), subducted

Nazca lithosphere is apparently undeformed by folding along strike and dips moderately eastward beneath the Maule rupture region (Hayes et al., 2012). South of 40°S, subduction zone seismicity decreases greatly, rendering estimation of the slab dip less precise; the Nazca plate subduction dip here has been estimated to be ~30° (Gutscher

2002) between 40-43°S. Just to the north of the Nazca-South America-Antarctic triple junction, where the active spreading Chile ridge subducts (~46°S), the Nazca slab subducts eastwards with a dip of ~35° (Russo et al., 2010a). An active volcanic arc is present from the Maule region to the Chile triple junction. In both the Peruvian and central Chilean flat-slab segments, the subhorizontal lithosphere of the Nazca slab extends for hundreds of kilometers eastwards before transitioning to a steeper dip

(Cahill and Isacks 1992; Anderson et al., 2005). A conclusive explanation for the segmented flattening of the Nazca slab remains elusive; the most likely mechanism invoked to account for flattening is based on the fact that the flat slab regions coincide with the projected locations of seamounts and aseismic ridges currently observed in the bathymetry of the Nazca plate – presumably, subduction of these thick, buoyant crustal features diminishes the negative buoyancy of the Nazca slab, resulting in the observed flat geometry (Cahill and Isacks 1992; Gutscher et al., 1999; Anderson et al., 2007;

Hayes et al., 2012).

1.3.4 Bathymetric Features

On the oceanic Nazca plate, there are a number of bathymetric features entering the subduction trench (e.g. hotspot tracks, fracture zones) that act to segment the

Chilean margin. The Valdivia Fracture Zone (VFZ) is a system of former transform faults that divides the Nazca oceanic lithosphere into two components: (i) the oceanic lithosphere to the north of the VFZ that formed at the Pacific-Farallon/Nazca spreading center > 20 Ma (Muller et al., 1998) and (ii) the oceanic lithosphere to the south of the

VFZ that was created at the Chile Rise < 20 Ma (Herron et al., 1981). Extending from the VFZ is the Mocha Fracture Zone (MFZ), another system of former transform faults.

Considered together, the VFZ, the MFZ, and the Chile Trench form a triangular area, the Mocha Block, which acts as a boundary between the two segments of the Chilean subduction zone (Barrientos and Ward, 1990).

Farther to the north, another prominent bathymetric feature on the subducting oceanic Nazca plate is the Juan Fernandez Ridge (JFR). The JFR runs E-W along the

Nazca plate and formed from the JFR hotspot (Yanez et al., 2001). The JFR is currently subducting at the Chile Trench ~33°S latitude (von Huene et al., 1997), near the

northern boundary of the Maule segment. The Nazca Ridge (NR) formed at the Easter

Island hotspot (Pilger, 1984) and is subducting obliquely to the Peru-Chile trench north of the study region.

1.3.5 Crustal Shortening

Despite the fairly constant convergence rate of the Nazca and South American plates since the Late Oligocene (~32-25 Ma) when the Farallon plate broke into the

Cocos and Nazca plates (Pardo-Casas and Molnar, 1987), there is remarkably variable shortening along the plate margin (Isacks 1988; Cahill and Isacks 1992; Russo and

Silver, 1996). Above both the Peruvian and the Argentina-Chile flat-slab segments, the

South American lithosphere is less deformed than crust along strike in the Arica bend region. The Andes experience the most pronounced along-strike lateral shortening and vertical thickening near the Bolivian orocline – Altiplano region (15-23°S) which has undergone more than 300 km of crustal shortening and thickening, associated with eastward-directed compressive stresses on the South American margin from subduction since the Cenozoic (e.g. Suarez et al., 1983; Kley et al., 1999; Elger et al.,

2005; Barnes and Ehlers, 2009).

1.3.6 Maule Rupture Zone

The complex evolution of the Chilean subduction zone is evident on a large-scale over the entire continental margin, and is even more obvious on a smaller scale when looking at localized surficial features. Here, we present a more detailed view of the region of central Chile affected by the Maule, 2010 earthquake (33°S-38°S), (see also

Section 1.2). The surface geology of the Andean Cordillera within the Maule rupture zone consists of three main strike-parallel morphostructural units: the Coastal

Cordillera, the Central Depression, and the Main Cordillera. These units trend sub-

parallel to the trench and can be attributed to subduction erosion and related migration of the trench arc system (Hussen et al., 2000; Willner et al., 2005, 2009); the Coastal

Cordillera and the Main Cordillera persist along the entire extent of the subduction zone in this region.

The Coastal Cordillera is the westernmost unit and incorporates the oldest magmatic arc (Late Paleozoic) in the study region. This arc is no more than 50 km wide

(Hartley et al., 2000) and is composed primarily of Jurassic-late Cretaceous basic to andesitic volcanics intruded by coeval plutonic belts (Giese et al., 1999; Cembrano et al., 2007) with a crustal thickness of ~45 km (Allmendinger et al., 1983). It can be subdivided into a Coastal Batholith and a metamorphic belt containing the Eastern and

Western Metamorphic Series (Charrier et al., 2007). The Eastern Series is composed of metagreywackes that were subject to low pressure/high temperature metamorphism related to intrusion of the Coastal Batholith in the Triassic (Willner et al., 2009). In contrast, the Western Series is an amalgamation of oceanic and continental rocks that were subject to metamorphism at high pressure/low temperature, including occurrences of blueschists (Willner et al., 2009).

The Central Depression is an elongated north-south trending basin that runs from

~33°S to ~46°S (Cembrano et al., 2007) and has variable width ranging from ~25 km to

~100 km (Hartley et al., 2000) with a continental Moho depth of ~30 km (Krawczyk et al., 2006). It is filled with ~1-2 km of continental, marine, and fluvial sediments of

Oligocene to Holocene age and formed as a result of extension in the late Oligocene – early Miocene (Glodny et al., 2006). Abundant lahar and volcanic avalanche deposits

also accumulated in the Central Depression, sourced from the Main Cordillera to the east (Charrier et al., 2007).

The Main Andean Cordillera is characterized by Mesozoic marine deposits

(Alvarado et al., 2005) and Mesozoic-Cenozoic volcanic and volcaniclastic units intruded by late Miocene-Pliocene plutonic rocks and the present-day volcanic arc

(Cembrano et al., 2007). Crustal thickness measurements vary in the Main Cordillera, with values ranging from 40 km (Krawczyk et al., 2006) to 65 km (Allmendinger et al.,

1983).

Figure 1-1. Historic earthquakes from Hayes et al. (2013) (a) Red ellipses show the rupture zones of large historic earthquakes along the Chilean margin (Beck et al. (1998)). Yellow ellipse outlines the rupture zone of the Maule event with the star as the epicenter. (b) Red triangles mark the locations of the IMAD seismic stations. Centroid moment tensor solutions for the main shock are given in the bottom left inset.

Figure 1-2. IRIS CHAMP temporary broadband seismic network. Red triangles represent broadband seismic stations that operated from March – October 2010.

Figure 1-3. CHAMP stations with alphanumeric labels. Each of the stations is identified as UXXB, with XX being equal to the numeric label shown on the black triangle in the figure.

Figure 1-4. From Kroner and Stern (2005), map of Gondwana at the end of the Neoproterozoic (~540 Ma) showing the arrangement of Pan-African belts. AS, Arabian Shield; Br, Brasiliano; DA, Damara; DM, Dom Feliciano; DR, Denman Darling; EW Ellsworth-Whitmore Mountains; GP, Gerlep; KB, Kaoko; MA, Mauretanides; MB, Mozambique Belt; BS, Nubian Shield; PM, Peterman Ranges, PB, Pryoiz Bay; PR, Pampean Ranges; PS, Paterson; QM, Queen Maud Land; RB, Rokelides; SD, Saldania; SG, Southern Granulite Terrane; TS, Trans-Sahara Belt; WB, West Congo; ZB, Zambezi.

Figure 1-5. Modified from Ramos (1988), terranes accreted to Southern South America and South American cratonic regions.

Figure 1-6. Modified from Hicks et al. (2014), the large-scale geologic units of the Maule rupture zone (33°S-38°S). Units shown include the Eastern and Western Metamorphic Series of the Coastal Cordillera, the Central Depression sedimentary basin, and the Main Andean Cordillera.

CHAPTER 2 SEISMIC ATTENUATION WITHIN THE MAULE, CHILE RUPTURE ZONE

2.1 Introduction

Seismologic, geodetic, and geologic studies generally indicate that the hallmark of both the mechanics of subduction and subduction zone structures is heterogeneity on scales considerably smaller than any individual subduction zone. Juxtaposition of diverse geology above and below the plate interface and along and across subduction strike combines with variable thermal and rheological structure to yield subduction zone segmentation and heterogeneous strain release of far field plate motion stresses on interplate subduction interfaces. The segment of the Nazca-South America subduction zone that ruptured during the Mw 8.8 February 27, 2010 Maule earthquake (hereafter termed the Maule segment) was the locus of geodetically observed interseismic strain accumulation deficit (Ruegg et al., 2009) and was subject to spatially variable interplate coupling prior to the 2010 earthquake (Lorito et al., 2011; Metois et al., 2012; Moreno et al., 2012), slipped heterogeneously during co-seismic rupture as observed seismically and geodetically (DeLouis et al., 2010; Lay et al., 2010; Moreno et al., 2010,2012; Lorito et al., 2011; Luttrell et al., 2011; Pollitz et al., 2011; Vigny et al., 2011), and generated heterogeneous post-seismic slip throughout the main shock rupture region (Vigny et al.,

2011; Lin et al., 2013; Luttrell et al., 2011). Furthermore, details of the rupture process, as revealed by back-projection of high frequency seismic energy recorded teleseismically, shows that the rupture itself proceeded heterogeneously along three adjacent portions of the Maule segment (Kiser and Ishii, 2011). The heterogeneity of interplate locking, co-seismic slip and geodetic strain release, and post-earthquake aseismic slip are consequences of the geologic and structural heterogeneity of the

Maule segment and segments of the Nazca-South America subduction zone adjacent to it.

Possible along-strike structures within the Nazca plate that might bound the

Maule segment include the Juan Fernandez Ridge or Challenger Fracture Zone, in the north, and the Mocha Fracture Zone (Contreras-Reyes and Carrizo, 2011) or the

Arauco Peninsula (Lin et al., 2013) in the south (Figure 2-1). The age of the Nazca plate lithosphere, subducting at ~6.8 cm/yr along the segment (Altamimi et al., 2007), varies apparently smoothly from around 28 Ma in the south where the Mocha FZ subducts to

~35 Ma in the north where the Challenger FZ enters the trench (Mueller et al., 1997).

The Nazca slab strikes approximately N-S at both the northern (34°S) and southern

(37°S) ends of the Maule earthquake rupture zone, as defined by the aftershock sequences (Hayes et al., 2013), but between these two areas, slab strike, as defined by the SLAB1.0 model of Hayes et al. (2012), is ~N30°E.

The offshore South America portion of the Maule segment is divisible into an active accretionary prism ~40 km wide, juxtaposed against an older, inactive accretionary complex extending landward to the Chilean coast (Moscoso et al., 2011).

Moscoso et al. (2011) correlated the transition from active to paleoaccretionary complex with the up-dip limit of seismic rupture during the 2010 Maule earthquake main shock: observed higher seismic velocities of the more highly metamorphosed older prism imply greater yield strength of these rocks where they are in contact with the subducted

Nazca plate, leading to seismogenesis rather than aseismic slip. Variable thickness of sediments in the subduction channel, either carried by the Nazca plate or eroded from the accretionary prism, can also affect seismic rupture segmentation, as a thicker

sediment-filled subduction channel may diminish the roughness of the Nazca slab surface, thereby allowing rupture to propagate across smoothed asperities (Contreras-

Reyes and Carrizo, 2011).

Above the interplate interface, sub-aerial South America is divisible into a

Coastal Cordillera, the Central Depression, and the Andean Cordillera (Figure 2-1). The

Coastal Cordillera is a composite of late Paleozoic-Mesozoic turbidites metamorphosed to greenschist facies and intruded and covered by several generations of Mesozoic arc- related plutons and volcanic extrusives (Willner 2005; Creixell et al., 2006). Sporadic occurrences of blueschist facies rocks and imbricates of basaltic oceanic crust (Willner

2005) indicate that the turbidites were originally part of Permian-Triassic subduction complex (Lucassen et al., 2004) that was subsequently active as an arc during the Late

Paleozoic (Willner, 2005).

East of the Coastal Cordillera, basement rocks are buried beneath the Central

Depression, a longitudinal valley with infill of Quaternary sediments, volcanic, and plutonic rocks (e.g. Charrier et al., 2007). To the north, the flat basin of the Central

Depression terminates at ~33.5°S, where the more modest topography of the Coastal

Cordillera is juxtaposed directly against the high elevations of the Andean Cordillera.

The Andean Cordillera to the east is the site of the currently active volcanic arc, built largely upon Cenozoic sediments, volcaniclastics, and intrusives (e.g. Charrier et al.,

2007).

In our study area, South American lithosphere is comprised of the Chilenia

Terrane (Ramos et al., 1986; Alvarado et al., 2007; Willner et al., 2008), which includes a Paleozoic Gondwanan (Domeier and Torsvik, 2014; Hyppolito et al., 2014) continental

crustal basement subsequently extensively intruded during a long history of subduction.

Crustal thicknesses in the region have been determined by both body wave tomography

(Haberland et al., 2009; Hicks et al., 2014) and analysis of receiver functions

(Dannowski et al., 2013). Dannowski et al. (2013) estimate the continental Moho lies at

~38 km depth at a latitude of ~36.5°S; Haberland et al. (2009) resolve the Moho at ~50 km depth at 37-38°S, and Hicks et al. (2014) detect upper mantle velocities at depths of

45-50 km beneath South American crust of much of the Maule segment. Waveform modeling of regional earthquakes indicates best fitting waveforms occur for crustal thicknesses between 30-45 km for the northeastern portion of the Maule rupture zone

(Alvarado et al., 2007). Variable crustal thickness of the overriding plate (compare

Haberland et al., 2009, and Dannowski et al., 2013) likely alters the locus of slab-mantle wedge contact, and therefore could also change the interplate contact area – with implications for the seismogenic zone (Hicks et al., 2014). Along strike changes in the down-dip position of the slab-mantle wedge contact may also contribute to segmentation of seismic rupture.

We assess the potential effects of structural heterogeneity that lead to megathrust rupture segmentation using local P and S waves recorded at the IRIS

CHAMP temporary seismic network to calculate the differential seismic attenuation along aftershock event-station ray paths. We then incorporated the path-integrated attenuation results into a three dimensional inversion that yielded the volumetric variation of seismic attenuation within and above the 2010 Maule, Chile rupture zone.

We assessed the quality of our model by examining the ray path hit count distribution in each model block, in addition to performing synthetic testing to see what size feature we

can resolve. We compare our attenuation model with previous studies in the Maule region and draw correlations with local geology (e.g. Vásquez and Franz, 2008;

Vásquez et al., 2011), P wave velocity structure (Hicks et al., 2012, 2014), and megathrust rupture segmentation (Kiser and Ishii, 2011).

2.2 Seismic Attenuation Background

Attenuation is the decay of a seismic wave’s amplitude as it propagates over an increasing volume and is quantified by the dimensionless quality factor, Q, as described in Appendix A. Theory and experimental observations show that attenuation is a function of temperature (e.g. Karato and Spetzler, 1990; Faul and Jackson, 2005), fluid content (e.g. Mavko and Nur, 1979; Karato 2003; Aizawa et al., 2008), and melt porosity

(e.g. Hammond and Humphreys, 2000; Faul and Jackson, 2007), and is therefore often used to characterize the Earth medium (e.g. Knopoff, 1964). Thermal variations strongly influence the attenuation of seismic waves as they propagate, with higher temperatures causing exponentially higher attenuation (e.g. Faul and Jackson, 2005). This direct correlation is so ubiquitous that there exists a quantitative relationship between temperature and attenuation (e.g. Faul and Jackson, 2005; Jackson et al., 2002). In

2008, Aizawa et al. conducted a pilot laboratory study to investigate the effects of water on upper mantle seismic attenuation and found that dunite samples containing ~0.3%

H20 were susceptible to attenuation. Prior to this study, the influence of water on seismic attenuation was based entirely on estimates, with no experimental support. Faul et al. (2004) reported an increase in attenuation from laboratory studies when small amounts of melt porosity (>1%) were introdued to mantle minerals.

Previous seismic attenuation studies have used local and regional seismic data

(e.g. Bowman 1988; Roth et al., 1999; Schurr et al., 2003; Stachnik et al., 2004; Russo

et al., 2005; Deshayes et al., 2006; Wiens et al., 2008; Pozgay et al., 2009) to characterize seismic attenuation in a variety of tectonic environments. In subduction zones, certain large-scale features such as subducting plates and sub-volcanic arc regions have been imaged unambiguously (e.g. Roth et al., 1999; Stachnik et al., 2004).

For the South American subduction zone in particular, four attenuation tomography studies have been conducted north of the Maule segment, all showing a high-Q east- dipping Nazca slab overlain by a zone of higher attenuation (Myers et al., 1998; Schurr et al., 2003; Deshayes et al., 2006; Liang et al., 2014). Myers et al. (1998) developed a joint velocity-attenuation model between 18°S and 22°S and attributed the moderately low attenuation values in the lithospheric mantle to be a result of limited infiltration of wedge volatiles and/or partial melt. Schurr et al. (2003) generated a Qp tomography model of the Bolivian Andes (~20°S-25°S) and explained that high attenuation under the active volcanic front and portions of the Altiplano-Puna plateau could be representative of high temperature zones and mid-crust partial melting. Just south, ~24°S-29°S, body wave attenuation tomography results reveal a west-dipping high attenuation zone beneath the Eastern Cordillera which Liang et al. (2014) interpret to be a piece of delaminated South American lithosphere. In Central Chile (~30°S-34°S), Deshayes et al. (2006) presented Qp and Qs measurements revealing a high S-wave attenuation beneath the volcanic belt, south of 33°S, in addition to high attenuation coincident with the location of the mantle wedge south of 33°S. North of this latitude, they did not observe the mantle wedge, likely due to flat slab subduction. Below 34°S, the southern extent of the Deshayes et al. (2006) model, there have been no studies investigating the seismic attenuation of the Chilean subduction zone.

2.3 Data and Methods

Our study uses data from a subset of IMAD, IRIS’ 58 station broadband seismic network, CHAMP, which operated from March 2010 through October 2010 (Figure 1-2 and Figure 1-3). We used local data and selected 197 suitable aftershocks with locations determined by Rietbrock et al. (2012), yielding 1,448 ray paths. We selected events that traveled primarily through the Nazca slab, mantle wedge, and overlying continental crust and avoided events that nucleated west of the forearc. The majority of the aftershocks that we used occurred either in the mantle wedge, on the subducting plate interface, or within the subducting plate, with our deepest event occurring at 126 km. We did not impose a magnitude threshold for the aftershocks and incorporated data with Mw ranging from 1.9 to 6.7. Because we used local events, we band-pass filtered our data between 0.7 and 20 Hz. We also applied a Hanning taper to each seismogram and removed the mean and trend from the seismograms.

We used local waveforms to calculate t*, which is related to Q through the travel time of the seismic wave as t*=t/Q (Schlotterbeck and Abers 2001), using two complementary spectral ratio methods (Russo et al., 2005). To calculate attenuation along an individual event-station ray path, these phase-pair methods estimate the differential attenuation between the compressional and shear waves and assume identical P and S travel paths for the same event-station pairs. By assuming identical travel paths, frequency domain division effectively removes the source-time function and the instrument response function, which we take to be identical for both P and S phases. This allows us to compare attenuation of arrivals with high signal to noise ratios via a ratio of spectral amplitudes of the two phases in varying frequency bands.

We visually identified the P and S phases on the vertical and transverse seismograms respectively (Figure 2-2). To assess the degree to which the measurements are affected by seismic noise, we also identified a band of pre-signal noise and compared the spectral content of the noise to that of the P and S waves. We only retained signals that have P and S spectra that are well above the ambient noise level. We then windowed each P and S wave arrival. Each window includes two, contiguous parts, the first containing the phase arrival and the second containing the arrival coda (Russo et al., 2005) (Figure 2-3). Typically, the entire P and S wave windows including both the signal and the coda are 1-2 seconds long. We note that in practice, reasonable variations in window length do not significantly affect the resultant estimate of attenuation for the phase pair (Bowman 1988, Roth et al., 1999, Russo et al., 2005). We divided the second portion of our phase window, that containing the phase coda, into 20 equal time segments. We invoked an evolving time window spectral calculation in which the leading portions of the P and S windows are included in all iterations, and each subsequent iteration incorporates an additional 1/20th of the trailing window, until the entire two-part window – phase arrival + coda – is used in the spectral ratio calculation. Thus, with each iteration, the length of the time window for which the spectral content is evaluated increases until all 20 subdivisions of the coda window are included in the calculation. Implementing this evolving window separately for each P and S wave yields 20 P wave windows and 20 S wave windows to be ratioed, producing

400 spectral ratios for each source-receiver ray path. For each of these 400 time series, the natural logarithm of the iteratively evolving spectra is calculated and plotted against frequency. We visually identified the segments of the frequency band for which the S

spectral amplitudes are greater than those of P, and for which both are above pre-signal noise, and calculated t* from the slope of the ratio within the identified frequency band.

This frequency band was no longer than 10 Hz in length, as unwanted scattered energy is often incorporated when longer frequency bands are used. Thus, each measurement yields 400 individual Qs/Qp spectral ratios. Because bulk attenuation values are often considered negligible (Pozgay et al., 2009), we ignore the contribution of bulk attenuation in this study and consider the measurements to be representative of Qs.

From these 400 individual ratios, we calculated an average Qs, and standard deviation.

We then determined a second estimate of Qs for the given source-receiver pair by stacking the spectral ratios produced for each of the 400 iterations, normalizing them, and taking the natural logarithm of the resulting composite spectral ratio, yielding a second Qs value (Russo et al., 2005) (Figure 2-4). This composite Qs measurement is in principle more robust than the 400 individual Qs measurements because the spectral stacking fills holes in the spectra and suppresses noise in the data. For this reason, we use the Qs values obtained from the composite stack method in our inversion for attenuation structure.

The first Qs measurement method is a very sensitive indicator of changes in attenuated energy of the arriving P and S energy and can be used to identify the arrival of multipathed arrivals, a phenomenon that occurs when seismic waves originate from a given source event and travel different paths to arrive at nearly the same time at a station (Russo et al., 2005). Because the arrivals traverse different structures on their way to the station, they may be attenuated differently; if all such phases are included in the attenuation measurement, the resulting Qs value is spurious and cannot reliably be

included in a tomographic inversion scheme for attenuation structure. Thus, we visually isolated multipathed phase arrivals by plotting the 400 Q values, noting systematic changes in the observed Qs that were clear functions of the measurement time windows, and recalculating Qs after excluding the portions of either the P or S (or both) windows corresponding to the onset of multipathed phase energy. We retained Qs measurements for use in the inversion for attenuation structure when the Qs value obtained from the stacked spectra was within the standard deviation of the 400 individual Qs measurements (Russo et al., 2005). Because these Qs measurements represent the path-averaged attenuation of the wave front accumulated along the entire event-station path, we implemented a 3D inversion to further constrain Qs within the rupture zone.

The t* values that we calculated using the spectral ratio method represent the path-integrated attenuation that the ray experiences from the hypocenter to the station.

To constrain the attenuation along the path, we can invert for Qs by relating t* to Qs along the ray paths through a velocity model. We tested two different velocity models.

We traced our rays through the IASPI91 radial-earth velocity model (Kennett 1991) and the 3-D velocity model of Ward et al. (2013), generated from ambient seismic noise tomography of the IMAD dataset. The two velocity models produced similar results and we present the results using the radial-earth model. We performed a three-dimensional cubic spline interpolation for each event-station ray path in order to gather samples at every 1 km depth. We divided the rupture zone into cubes and assigned constant velocity to each. The model is divided into 15 longitudinal blocks, 37.6 km long, and 25

latitudinal blocks, 40 km long. We inverted for structure from the Earth’s surface to 140 km depth; the depth dimension of each block was 9.33 km.

The observed vector of differential t* values is expressed as, M é ù *obs *calc -1 t p t - t = Q t - (2-1) ij ij å( s )i ê s ú i=1 ë g û ij

*obs where tij represents the t* value retained from the spectral ratio calculation for ray

*calc t s path j and model parameter (i.e. tomography cube) i, ij is the t* predicted value, Q is the value of shear attenuation in each grid block, tp and ts are the travel times of P and S waves, respectively, of ray path j in model parameter block i, and  = (Qp/Qs) and relates the attenuation of the compressional and shear waves. We tested a number of different  values ranging from 0.7-2.6 based on estimates from previous studies (e.g.

Roth et al., 1999; Shito and Shibutan, 2003a; Stacknik et al., 2004; Pozgay et al., 2009) and found that the optimal  value to minimize data misfit and model residuals is 2.0.

The generic inverse problem formulation given in Eq. 2-2 can be expressed for seismic attenuation as Eq. 2-3,

di  Gijmj (2-2)

1 (2-3) t *N  TNM Qs M where M is the number of blocks, N is the number of ray paths, t* = t*obs - t*calc, and

-1 T=[ts-tp( )]. This form applies for inversions of attenuation structure in one to three dimensions and requires no a priori estimates of the attenuation structure. We used a starting model of a constant 1/Qs of zero. We also used a second derivative smoothing

constraint in the inversion. The travel-time kernel is augmented with a weighted version of the smoothing matrix, S, and the data vector is padded with N zeros,

t * T  1  Q (2-4)     s   0  wS where the G matrix is now N+D rows by M columns.

To minimize the station response we also add a site damping term, , to the G matrix. To incorporate this, we horizontally concatenated a station response matrix, R

(size (N+D) x (NSTATIONS)) to the G matrix, containing ones where there is a matching event-station pair and zeroes elsewhere. We pad the rest of the G matrix with zeroes, and add the  term to the bottom of the columns with positive event-station pair correlations. We pad the data matrix with another zero, and in matrix form this now becomes

(2-5)

where w is a dimensionless weighting factor for smoothing and R incorporates ones and zeroes for the station-event pairs. Additionally, from this inversion we can determine the site response terms for each station, represented by Z in the model vector. The dimensions of the data vector are [(N+D+M)x1], the dimensions of the final G matrix are

[(N+D+1)x(M+NSTATIONS)], and the dimensions of the model vector are

[(M+NSTATIONS)x1]. To select the optimal smoothing value and site damping term, we inspected the data variances and output models deriving from different w and  values

and found that using a w of ~1000 and  of ~0.5 (Figure 2-7) will smooth large poorly constrained variations between adjacent blocks while allowing robust small-scale structures to remain intact.

Our initial unbounded inversion produced some blocks with non-physical results, i.e. Q<0. To avoid such results, we implemented a bounded linear inequality inversion with Qmin=0 and Qmax=1200. We used the implementation described by Menke (1989) to minimize ||d-Gm||2 subject to the constraints Hm ≥ h where m=m+m0, h=(1/Qmax) and H is the identity matrix. Although the bounded linear inequality inversion yields a model with some extremely high Qs values, reaching our Qmax, the non-negativity constraint eliminates models that are physically unreal. We removed outliers beyond 3 and re- ran the inversion excluding those data.

2.4 Results

Our 3D attenuation tomography model for the Maule rupture zone is shown as trench-normal latitudinal sections in Figures 2-9 through 2-14 and as depth sections in

Figures 2-14 through 2-20. On a large-scale, the inversion shows the zones of highest attenuation (Qs~200) at the surface, with attenuation decreasing (i.e., greater Qs values) progressively with depth. We observe a continuous east-dipping ~30° low attenuation feature (Qs > 800) persisting through all cross-sections of the model volume extending from the surface to the maximum model depth, 140 km. In Figures 2-9, 2-10, and 2-14, above this dipping, low attenuation feature Qs values rapidly decrease from Qs~800 to

Qs~250. Figures 2-11, 2-12, and 2-13 also show higher attenuation above the dipping feature but do not exhibit the abrupt change we observe in the other cross-sections.

Near the shallow forearc, we observe high Qs. Within the ~50 km of the continental

crust, we observe regions of higher attenuation north of ~34.5°S and south of ~37°S, punctuated by zones of lower attenuation between these latitudes. In Figure 2.11, the model shows a large, ~50 km wide, region of low attenuation above ~20 km depth, residing in the continental crust near 35°S, 71.5°W. This feature appears to interrupt the otherwise continuous attenuation distribution we see at this depth and latitude. Figure

2.12 has a similarly low attenuation feature in the crust along the eastern edge of our model volume, ~60 km wide. This body appears to extend to only ~15-20 km depth, shallower than the northern section, and is located at 35.8°S, 71.2°W. Perhaps the most outstanding feature in this model is the large low-attenuation body in Figure 2.13 that appears to be continuous with the eastward low-attenuation dipping body. This upward protruding feature appears to be ~50 km wide and ~20 km thick and is also evident in all depth sections (Figures 2-15 – 2-20), appearing at ~36.5°, 72.5°W.

2.5 Model Resolution

The ray path hit count distribution throughout the rupture zone clearly shows that the best resolution is in the center of the CHAMP network, ~35°S. We also have the best resolution within the top 50 km of the model volume; below this depth the ray paths sampling becomes much more sparse (Figures 2-21 - 2-25).

Our goal is to interpret well-resolved attenuation structure in light of the subduction zone tectonic processes outline above. Therefore, we conducted resolution tests in which we inverted synthetic data to see how well we could recover an original input structure. We designed a checkerboard input structure of alternating Q values of

200 and 500 with each cube having dimensions of ~188 km in latitude, ~200 km in longitude and ~9 km in depth. We calculated synthetic t* values for the actual ray paths of the observed event-station pairs, added normally distributed random noise with the

same data variance as errors from our actual t* data, and inverted to obtain checkerboard resolution images.

Horizontal section (i.e. ‘maps’ at depth) resolution tests reveal that the greatest resolution lies towards the shallower depths of our model volume (depth < 50 km).

Distinct blocks are resolvable in the center of the rupture area, with edge effects causing higher attenuation regions to be pushed towards the edge of the model. At the shallowest horizontal sections of ~5 km at 14 km depth, we are able to clearly resolve the checkerboard pattern in the center of the rupture zone. The resolution at deeper horizontal sections begins to diminish as we have less ray path coverage. We have the greatest vertical resolution in the center of the rupture zone (~34°S- ~36°S); however the vertical cross-section resolution tests show smearing in the vertical direction. This is a consequence of the source event-station sampling, which entails almost entirely near vertically incident ray paths (Figures 2-26 and 2-27).

2.6 Discussion

2.6.1 Large-Scale Subduction Structures

The attenuation structures imaged in this study are, on a large-scale, very similar to structures found in many other subduction zones (e.g. Schurr et al., 2003, Stachnik et al., 2004; Pozgay et al., 2009). The Qs values we determined are consistent with several other studies of seismic attenuation in subduction zones (e.g. Roth et al., 1999;

Bowman, 1998; McNamara, 2000; Stachnik et al., 2004; Rychert et al., 2008), but direct comparison of Qs measurements across tectonic environments is difficult as different data processing methods, such as filtering or inclusion/exclusion of bulk attenuation affects, may influence the results. Instead, we focus on the relative difference of Qs throughout our model. We clearly image an east dipping ~30° low-attenuation feature

that we interpret to be the subducting Nazca oceanic slab. The depth to the top of the slab coincides with the location of the slab upper surface as defined in the SLAB1.0 model (Hayes et al., 2012), thus allowing a calculation of the mean Qs both above and within the slab. The average Qs of the slab is ~870 and the region above the slab is

~505. The low attenuation relative to the remainder of the mantle wedge that we observe in the shallow forearc is a phenomenon that has been observed in other subduction zones including Central America (Rychert et al., 2008), the Bolivian Andes

(Schurr et al., 2003), and Alaska (Stachnik et al., 2004). Other studies have attributed this low-Qs to a cold mantle wedge nose that is separated from the large-scale flow of hot materials in the remainder of the wedge. This may be the case for the Maule segment as well, although ray path coverage in this region of the model is not optimal and, as a result, suggests that these results may be poorly constrained.

The inversion technique we used is susceptible to producing spurious structure, i.e., edge effects, and could possibly yield high Qs blocks along the edges of our model.

However, we do not observe high Qs values on all sides of the model, as would be expected for spurious structure. Instead, high Qs values appear along only the western edge of the model, coincident with the location of the shallow Nazca slab. The lack of a high Qs edge at the top of our model (i.e., along the surface) suggests that we resolved real structures.

2.6.2 High-Qs Bodies at the Surface

We see significant latitudinal variation of Qs at the surface of the rupture zone

(depth < 20 km). A correlation between seismic attenuation and surface geology may be observed in our inversion if the model block sizes are small enough to resolve surface features. With block sizes of 9 km vertically, we cannot make any conclusions regarding

structures smaller than this dimension. The two low attenuation bodies in Figures 2-11 and 2-12 are both larger than 9 km, and we can confidently say that we are resolving real features. These two bodies are spatially coincident with dense, ultramafic Triassic intrusions (Vásquez et al., 2011), which would be manifest in an attenuation tomography as high Qs bodies. The correspondence between the Triassic plutons and the locations of the low attenuation bodies suggests that we are resolving the plutons in our tomography model.

2.6.3 Cobquecura Anomaly

In Figure 2-13, we image a low attenuation structure (Qs~800) approximately 20 km thick and 50 km wide apparently protruding upwards from the subducting slab. This high Qs structure corresponds to the Cobquecura anomaly (CA) found in the velocity tomography model of Hicks et al. (2014), an anomalous body with Vp > 7.5 km/s, representing a P wave velocity increase of 8% relative to surrounding South America crust (Figure 2-28, Figure 2-29). Hicks et al. (2014) provide a detailed discussion of this high velocity anomaly and its correlation with the Bouguer gravity field, models of pre- seismic locking, co-seismic slip models, high frequency energy propagation, aftershock seismicity distribution, and surface geology. Based on their calculations of the regional

Bouguer gravity, Hicks et al. (2014) found a strong positive gravity anomaly (-70 mGal) at the surface just above the location of the CA (Figure 2-30). They also note that the

CA is coincident with high pre-seismic locking, small co-seismic slip, a step down-dip in rupture propagation (Figure 2-31), and reduced aftershock activity. Additionally, the occurrence of Triassic plutons punctuating the Coastal Cordillera at latitudes north of

37.5°S suggests that similar plutons may exist south of this latitude but may be seated within the continental crust and not exposed at the surface. Outcrops with the requisite

petrologic composition occur in the vicinity of the seismic CA: the Triassic Cobquecura pluton is composed of gabbroic rocks and mingled gabbro and olivine bearing fayalite- granodiorites which are estimate to be comagmatic, crystallizing between ~210-203 Ma

(Vásquez and Franz, 2008; Vásquez et al., 2009; Vásquez et al., 2011). Based on these data, Hicks et al. (2014) suggest that the CA is a dense, ultramafic body emplaced during Triassic extension through magmatic processes (e.g. asthenospheric upwelling, slab steepening, or slab detachment), coeval with the Triassic intrusives observable on the surface north of 37.5°S, particularly the exposed Cobquecura pluton. Given the ~15 km of seismically different velocities (relatively slow, appropriate for typical continental crust) currently lying between the anomaly and the pluton, a co-genetic origin for the two structures would imply their separation at time of genesis due to structural and/or magmatic processes, or later separation and exhumation of the Cobquecura pluton from the currently deep crustal seismic structure.

While we agree that the CA is likely characterized by a dense, ultramafic body, we suggest a different emplacement mechanism than that proposed by Hicks et al.

(2014). To achieve ultramafic compositions in the continental crust, magmatic processes alone are not sufficient to melt the required amounts of peridotite to generate a komatiite melt (i.e., ultramafic volcanic origin). Instead, a more plausible mechanism would be a tectonic process in which a rifting episode formed a small oceanic basin - with an active spreading center - that subsequently closed, with the spreading center either obducting or subducting to shallow depths and then imbricating or accreting to the overriding South American crust. The composition of the nearby Cobquecura pluton suggests that it developed in an anhydrous tectonic environment (thus allowing olivine

bearing granitoids to form mingled with gabbros), indicating an extensional regime rather than a subduction environment (Vásquez et al., 2011).

Hicks et al. (2014) suggested that the CA played a role in controlling the distribution of slip during and after the 2010 rupture, consistent with results from the back-projection of high frequency seismic energy recorded in Japan and the USA which revealed that high frequency seismic radiated from the 2010 Chile megathrust propagated mostly SW-NE but stepped down-dip around the CA ~36°S (Figure 2-31)

(Kiser and Ishii, 2011). Kiser and Ishii (2011) relate heterogeneity of rupture propagation

~36°S to historical seismicity, and note that this region corresponds to the northern extent of the estimated 1835 rupture zone and the southern extent of the 1985 rupture zone, suggesting that a rupture blocker may exist at this latitude. While the CA may have served as a barrier to rupture in the past, which would be consistent with historic rupture zone distributions, it was not able to arrest the high frequency energy radiated from the Maule megathrust. When we consider that the Maule event epicenter occurred adjacent to the CA on the subducting plate interface, it is plausible that the CA may have incited the nucleation of the megathrust. Models of pre-seismic locking (e.g.

Moreno et al., 2010, 2012) and interseismic locking (e.g. Metois et al., 2012; Moreno et al., 2010, 2012) support this argument, and clearly show a high degree of locking on the plate interface just west of the CA. This is consistent with the argument proposed by

Ruff (1992) that asperities in subduction zones influence the rupture process and heterogeneity for large earthquakes, and with observations of Tassara (2010) that other ruptures along the Central and Southern Andean margin nucleate near asperities in the forearc.

2.7 Conclusion

We presented a 3D seismic attenuation tomography of the rupture zone of the

Mw 8.8 Maule, Chile earthquake using local aftershock data recorded within the IRIS

CHAMP temporary seismic network. Attenuation measurements were made using a spectral ratio phase-pair method (Russo et al., 2005) in which we iteratively calculated

Qs for an evolving time window incorporating both the phase signal and coda. We used these path-averaged attenuation values (Qs) along with the IASPEI radial earth velocity model in a non-negative linear least squares inversion to solve for an attenuation model.

Tomography results clearly show a subducting slab (Qs~870) with low attenuation that is overlain by a highly attenuating mantle wedge and South American crust (Qs~505).

These large-scale features are consistent with the thermal regime present in subduction zones. We observed smaller, low attenuation features that punctuate the otherwise continuous high attenuation of the South American continental crust surface at depths less than 20 km. These bodies are likely representative of Triassic intrusives. Perhaps the most notable feature of the attenuation model is a high Qs structure that appears to be attached to the subducting slab and that is spatially coincident with, and approximately the same size as, the Cobquecura velocity anomaly discovered by Hicks et al. (2012, 2014) in a high resolution P wave tomography study. Our characterization of this low attenuation feature is consistent with that suggested by Hicks et al. (2014) in that it is ultramafic in composition, and we suggest that the CA developed in a dry, hot marginal basin spreading center extensional regime during the Triassic and was subsequently obducted or subducted to shallower depths and accreted to the overriding

South American crust. The proximity of the highly-locked CA to the nucleation site of the

Maule event strongly suggests that there is a relationship between the interplate locking

of asperities with nucleation of megathrust earthquakes in subduction zones.

Additionally, the high frequency seismic energy propagation following the 2010 Chile earthquake steps down-dip around the low attenuation feature, suggesting that this body may influence rupture segmentation or act as a barrier to rupture. Overall, this attenuation tomography of Central Chile shows large-scale attenuation structures that are consistent with attenuation in other subduction systems, but also effectively resolves smaller scale features that we interpret in terms of regional properties and that have implications for along-strike rupture heterogeneity and seismogenesis.

Figure 2-1. Main morphostructural units in study area.

Figure 2-2. The P, S, and pre-signal noise windows are shown to demonstrate the evolving time window method for the April 5, 2010 seismic recording at CHAMP station U15B. The 0 marker indicates the event origin time. The P wave window is highlighted in red and the S wave window is highlighted in blue. The preceding orange window contains the portion of the pre-signal noise used for spectral analysis comparison with the main P and S phase spectra.

Figure 2-3. Illustration of the evolving time window method. (Top) P wave window (Bottom) S wave window. The first portion of the windows are included in all iteration evaluations of the spectra and the second portions are subsequently added in increments of 1/20th the trailing coda window to the spectral analysis using an evolving window technique.

Figure 2-4. Frequency spectra for event shown in Figure 2-2 and Figure 2-3. (Top left) 1 of 20 spectrum for P phase and (top right) S phase for the April 5, 2010 aftershock contained in Fig 2-2 and Fig 2-3. The solid line represents the signal and the dotted line represents the noise spectrum. The X axis is frequency in Hertz and the Y axis is the amplitude spectrum. (Bottom) The stacked spectral ratio for the April 5, 2010 event yields a more linear relationship with frequency. The solid line between ~6 Hz – 14 Hz denotes the frequency band for which this event was evaluated.

Figure 2-5. Illustration of the ray path distribution in the Maule rupture volume. Colors indicate different Q values for the respective ray path.

Figure 2-6. Illustration of the tomography concept. Each ray path traverses many model parameters (i.e., cubes) but yields a single t* value at the recording station. We use the intersection of ray paths found in a single cube to constrain t* along a given ray path (j) in a particular cube (i) and, with the IASPEI velocity model, we can invert for Qs in that cube. The pink cube in the figure contains segments of all four ray paths shown.

Figure 2-7. Plot of data misfit versus model residuals for different damping values tested in the inversion process. This ‘L’ curve shows that the optimal value for damping to minimize both the data misfit and the model residuals is ~0.5.

Figure 2-8. Study area for the Maule attenuation tomography. White circles are the aftershocks used. Red triangles represent Quaternary volcanoes. The 6 slices shown through the rupture zone (A-A’, B-B’, C-C’, D-D’, E-E’, F-F’) correspond to slices through the attenuation tomography model that are shown in the following figures (Fig 2-9, Fig 2-10, Fig 2-11, Fig 2-12, Fig 2-13, Fig 2-14)

Figure 2-9. Cross section A-A’ through attenuation tomography model. See Fig 2-8 for location of A-A’ slice. Black dashed line corresponds to the depth to the top of the Nazca slab determined from SLAB1.0 (Hayes et al., 2012). White circles represent aftershocks used in this region. Red triangles at the surface represent Quaternary volcanoes.

Figure 2-10. Cross section B-B’ through attenuation tomography model. See Fig 2-8 for location of B-B’ slice. Symbols are the same as Figure 2-9.

Figure 2-11. Cross section C-C’ through attenuation tomography model. See Fig 2-8 for location of C-C’ slice. Symbols are the same as Figure 2-9.

Figure 2-12. Cross section D-D’ through attenuation tomography model. See Fig 2-8 for location of D-D’ slice. Symbols are the same as Figure 2-9.

Figure 2-13. Cross section E-E’ through attenuation tomography model. See Fig 2-8 for location of E-E’ slice. Symbols are the same as Figure 2-9.

Figure 2-14. Cross section F-F’ through attenuation tomography model. See Fig 2-8 for location of F-F’ slice. Symbols are the same as Figure 2-9.

Figure 2-15. Depth slice at 5 km.

Figure 2-16. Depth slice at 14 km.

Figure 2-17. Depth slice at 23 km.

Figure 2-18. Depth slice at 33 km.

Figure 2-19. Depth slice at 42 km.

Figure 2-20. Depth slice at 51 km.

Figure 2-21. Ray path hit count for depth 0 – 9.3 km.

Figure 2-22. Ray path hit count for depth 9.3 – 18.7 km.

Figure 2-23. Ray path hit count for depth 18.7 – 28.0 km

Figure 2-24. Ray path hit count for depth 28.0 – 37.3 km.

Figure 2-25. Ray path hit count for depth 37.3 – 46.7 km.

Figure 2-26. Input Q model (checkered Q=200 and Q=500) and output model after resolution testing at the two shallowest, and best resolved, depth slices of 5 km and 14 km.

Figure 2-27. Input Q model (checkered Q=200 and Q=500) and output model after resolution testing at -34.1°S, -35.1°S, and -36.2°S latitude slices. Our best resolution is at the center of the CHAMP network, around 35°S. We observe vertical smearing because our dataset consists primarily of nearly vertically incident ray paths.

Figure 2-28. Cross sections used in Hicks et al. (2014) velocity model. Sections are labeled A through F from top to bottom. Figure 2-29 shows center cross section, cross section C-C’. White circles represent aftershocks used in this study and purple star represents epicenter of Maule Mw 8.8 megathrust event.

Figure 2-29. Comparison of attenuation tomography model to velocity tomography model. (Top) Hicks et al. (2014) P wave velocity tomography model. (Bottom) this study’s attenuation tomography model.

Figure 2-30. Comparison of attenuation tomography model to velocity tomography model and gravity. (Top left) Hicks et al. (2014) P wave velocity tomography model at 25 km depth. (Top right) Surface gravity signature. (Bottom) this study’s attenuation tomography model at 23 km depth.

Figure 2-31. Attenuation model at 23 km depth with high frequency rupture propagation dots from Kiser and Ishii (2011) superimposed.

CHAPTER 3 SHEAR WAVE SPLITTING OF TELESEISMIC EVENTS

3.1 Background

Whether minerals within the earth’s mantle are anisotropic or isotropic, in aggregate, is of particular interest in mantle dynamics, tectonics, and structural geology and can help us understand motions associated with plate tectonics such as mantle flow direction, in addition to the processes of internal plate deformation (Silver and Chan

1988). Shear wave splitting is a powerful tool in understanding the distribution of anisotropy in a particular area and is described in more detail in Appendix B. On a large-scale, because the a-type olivine fast axis in the upper mantle aggregates aligns in the direction of maximum strain (e.g. Nicolas and Christensen, 1987; Kaminski and

Ribe 2002), fast directions determined from shear wave splitting measurements are often interpreted in terms of mantle flow field direction. In this chapter, the data gathered from the teleseismic events recorded at stations within the Maule rupture zone provide an opportunity to analyze regional upper mantle properties. The results of the shear wave splitting measurements will be used to illuminate mantle flow direction below and around the subducting Nazca slab.

3.2 Previous Teleseismic Shear Wave Splitting Studies in South America

The South American subduction zone has been the site of many shear wave splitting studies, in an effort to better constrain the mechanisms responsible for the observed shear wave splitting. In 1994, Russo and Silver published a seminal paper in shear wave splitting based on their observations of fast polarization directions along the

South American subduction zone from 10°N - 35°S. They observed primarily trench parallel splitting directions in regions of normal subduction, and argued against the 2-D

model of slab-entrained mantle flow that was commonly assumed – largely on the basis of geodynamic models that were computationally limited to two dimensions (i.e. across strike) – in subduction zone settings (Ribe, 1989). Russo and Silver (1994) instead proposed that the presence of trench-parallel fast directions suggests that the sub-slab mantle is largely decoupled from the subducting Nazca plate and that the retrograde motion of the subducting lithosphere, as it retreats from the overriding South American plate, may induce sub-slab mantle flow with a strong trench-parallel component due to a gradient in normal stress that pushes mantle material laterally around the slab, i.e. stagnation point flow (Figure 3-1). Although the majority of their fast directions were trench parallel, Russo and Silver (1994) did observe deviations from this in their splitting dataset. To explain these deviations, they postulated that variations in slab morphology, i.e., transitions from normal to flat subduction, observed in several regions along strike, could explain the deviations from fast directions seen along strike of the South

American subduction zone as a result of localized pressure gradients in the mantle flow field. For example, localized pressure gradients resulting from changes in slab morphology near the flat-slab segments of the subduction zone could result in trench- normal splitting measurements. Following the study of Russo and Silver (1994), subsequent shear wave splitting studies in South America observed trench-normal splitting at the stagnation point area of normal slab subduction in the Peru-Chile borderland area (e.g. Polet et al. 2000) and trench-parallel splitting in other regions of the subduction zone (e.g. Anderson et al., 2004, MacDougall et al., 2012) and often invoked the retrograde flow model of Russo and Silver (1994) to explain these observations.

Bock et al. (1998) calculated shear wave splitting measurements using teleseismic data in the South American subduction zone between 12°S – 28°S. They observed that the majority of spitting directions paralleled the APM of the subducting

Nazca slab (N80°E) but noted a subset of measurements that paralleled the strike of the trench. They attributed their observations to slab-entrained mantle flow coupled to the subducting Nazca plate, with a secondary regime of anisotropy resulting from local deviations in slab morphology possibly responsible for the trench parallel fast directions.

Polet et al. (2000) studied the shear wave splitting in the South America subduction zone around ~20°S with observations from the Broadband Andean Joint (BANJO) experiment, Seismic Exploration of the Deep Altiplano (SEDA), and Projecto de

Investigacion Sismologica de la Cordillera Occidental (PISCO) seismic networks. They observed mostly E-W fast directions at the BANJO stations, but saw nearly trench parallel splitting at the SEDA and PISCO networks. They postulated that these observations were a result of a transition in mantle flow field regimes beneath the

Andean mantle and the Brazilian craton. In 2004, Anderson et al. measured shear wave splitting between 30°S-36°S using teleseismic data. They observed fast directions that were mostly trench-parallel in the southern and northwestern portions of their study area, but noted a gradual transition to more trench normal fast direction in the northeastern portion of their study area. Anderson et al. (2004) attribute this change in fast direction to the change in morphology of the subducting Nazca slab ~32°S, consistent with the retrograde flow model of Russo and Siler (1994). MacDougall et al.

(2012) made shear wave splitting measurements in the Maule rupture zone (33°S-38°S) and observed teleseismic fast directions that were predominantly trench-parallel.

Farther south, between 37°S -39°S, Hicks et al. (2012) found teleseismic shear wave splitting, recorded from the active source experiment, The Incoming Plate to Mega-

Thrust Earthquake Processes (TIPTEQ), with directions oriented with a dominant N-E fast direction consistent with the APM of the Nazca plate. This is not consistent with results from Russo et al. (2010a, 2010b) that show primarily trench parallel fast directions north of the Patagonian slab window.

While it is likely that many different mechanisms are responsible for localized mantle flow along different segments of the Chile subduction zone (e.g. rollback subduction, along-strike variation in slab dip and curvature, lateral heterogeneities in plate thickness, small-scale convection), for mantle flow fields that persist for hundreds of kilometers and are close to the subduction trench, the simple retrograde flow model of Russo and Silver (1994) is generally consistent with observations in other regions of the South America subduction zone. Results from more detailed studies, however, could reveal information about the smaller scale perturbations from this model and provide insight into the regional mantle flow field and variations in slab morphology.

3.3 Data and Methods

We used data collected from the IRIS CHAMP temporary seismic network comprised of 58 seismic stations, and selected 57 suitable teleseismic events (Table 3-

1) that occurred within 87° <  < 150° from the recording seismograph. After data analysis, we obtained measurements for 52 of the IRIS CHAMP stations. Of the potential 2,964 event-station pairs, we only used 386 event-station pair measurements, retaining only those with high signal to noise ratios after filtering between 0.01 and 10

Hz. We used the teleseismic SKS and SKKS phases to make splitting observations

because these waves travel through the outer core of the Earth. As a result, we can assume that these waves were radially polarized when they left the liquid outer core and that any source-side splitting the wave may have encountered was effectively erased.

Typically applied shear wave splitting methodologies of Silver and Chan (1991) and Wolfe and Silver (1998) use a grid search which seeks the splitting parameters,  and t, that best minimize the eigenvalues of the covariance matrix between the two horizontal seismogram components. These methods, applied to the Maule data set described above by Russo in 2010 (2013, personal communication) resulted in comprehensively poor measurements that precluded interpretation. We note that

MacDougall et al. (2012) were also unable to derive useful measurements from the

SK(K)S data set recorded at the CHAMP stations. Reasons for these failures stem from a combination of very short deployment time (6 months) – and therefore an insufficient number of suitable earthquakes generating high signal-to-noise SK(K)S phases – and a tendency for the Silver and Chan (1991) and Wolfe and Silver (1998) methods to underestimate uncertainties – leading to gross scatter among apparently acceptable results – for small data sets.

Therefore, we used a multi-channel analysis approach developed by Chevrot

(2000) that takes advantage of the fact that, for typical Earth anisotropy, the transverse component is simply the time derivative of the radial component multiplied by a factor that is dependent on both the time delay (t) and the angle to the fast axis (). This factor is termed the ‘splitting intensity’. If t is small compared to the dominant period of the signal, we can approximate Eq. 3-1 and Eq. 3-2 as Eq. 3-3 and Eq. 3-4 for a vertically incident shear wave traveling through a single horizontal layer of transverse

anisotropy where  is the angle between the fast symmetry axis and the initial polarization direction and (t) is the incoming wavelet.

 t  2  t  2 R(t)  t  cos  t  sin  (3-1)  2   2 

1   t   t  T(t)   t   t  sin 2 (3-2) 2   2   2 

R(t) (t) (3-3)

1 T(t)   t sin 2*(t) (3-4) 2

Thus, with SI=t sin(2), we see that the delay time (t) is related to the amplitude of the sinusoid and the angle to the fast axis ( in the equations) – i.e., the radial axis in teleseismic studies incorporating SK(K)S phases – can be determined from the phase of the sinusoid. Rewriting these simplified equations in vector form Eq. 3-5, we see a simple relationship between the matrix of transverse, or horizontal, component records to the derivative of the radial, or orthogonal horizontal component, through the splitting vector, s.

T  as r (3-5)

This suggests that the singular value decomposition of the data matrix provides a simple means to extract a single pair of splitting parameters for the full suite of SK(K)S data recorded at a given station.

To implement this method, we isolated the SKS or SKKS phase on the transverse and radial seismograms, deconvolved the radial and transverse components by the radial component, and projected the time derivative of the radial component onto the transverse component to calculate the coefficient of splitting (see Figure 3-3 and

Figure 3-4). For event-station pairs that had low RMS values, we retained the coefficient of splitting and azimuthal direction. We plotted the azimuth versus coefficient of splitting to fit a sin(2) curve to the data for multiple events recorded at a single station. From this sinusoidal fit, when considering all the data for a given station simultaneously, we were able to calculate the splitting intensity and splitting parameters  and t, for the station (see Figure 3-5).

An advantage to this multi-channel method is that most of the good quality data is useable and can be incorporated into fitting a splitting function. For example, measurements that would otherwise be considered null, i.e. when the polarization of the fast or slow axis of anisotropy is in the same direction as the initial wave polarization, by the grid search methods of Silver and Chan (1991) and Wolfe and Silver (1998) can be used in the averaging scheme of the Chevrot (2000) method. Additionally, the multichannel method provides more robust measurements with smaller uncertainties for data obtained from noisy environments than do the grid search techniques (Chevrot, 2000).

The main limitation to the multi-channel approach is the dependence on good azimuthal distribution of seismicity. Chevrot (2000) assessed the global distribution of earthquakes and found that, at teleseismic distances, this multichannel method is appropriate and optimal for stations in the South American subduction zone.

3.4 Results

The splitting intensity as a function of back azimuth is measured by projecting each transverse component record on the derivative of the radial component. From the splitting function, we are able to extract the splitting parameters  and t. The results are displayed in Table 3-2. We observe a strong component of trench parallel splitting with few, if any, trench normal fast directions observed at any station. Also, the fast directions observed do not parallel the APM of the subducting Nazca plate, but instead are dominated by splits rotated about 40-50° counterclockwise from Nazca APM. We observe delay times ranging from 0.3 s – 2.8 s and most stations, with two larger delay times at CHAMP stations U45B (-34.74°S, -71.64°W) and U62B (-35.66°S, -71.77°W) of

3.45 s and 3.46 s respectively. There is no consistent spatial distribution of delay times.

Large delay times are mostly found at stations sited in the Central Depression sedimentary basin, although some are coincident with the locations of batholiths in the

Coastal Cordillera. The majority of the smaller delay times are located along the entire length of the Chilean coastline, closest to the subduction trench.

3.5 Discussion

3.5.1 Time Delay

Shear wave splitting measurements are frequency dependent (e.g. Silver and

Savage, 1994; Marson-Pidgeon and Savage, 1997), with larger delay times obtained in studies that incorporate more low frequency energy and smaller delay times resulting from studies that include higher frequency energy. The band-pass filter that we implemented affected the wavelengths of the seismic records and, consequently, the time delays we obtained. Therefore, while the delay times that we found are generally consistent with typical determinations for shear wave splitting delay times of 0 – 2 s for

teleseismic phases subject to upper mantle anisotropy (e.g. Mainprice and Silver 1992;

Gledhill and Gubbins 1996; Kendall and Silver, 1996), it is difficult to directly compare our delay time measurements to delay times observed in other studies.

3.5.2 Contributions from Crustal Anisotropy

Splitting observations of core-refracted phases SK(K)S used in teleseismic shear wave splitting studies are reflective of receiver-side path-integrated splitting, making it difficult to determine the depth to anisotropy; this is an inherent limitation of the method.

However, the fast directions we observe are unlikely to be due to anisotropy in the crust, because the wavelengths of the SKS and SKKS phases are ~40 km and would not be very sensitive to features smaller than that dimension. Crustal thickness measurements in the Maule rupture zone do not exceed 50 km, with upper mantle velocities present at depths no greater than 45-50 km (Dannowski et al., 2013; Hicks et al., 2014) beneath the Maule segment. Additionally, the splitting times obtained at the majority of CHAMP stations are greater than 0.6 s, which is generally considered too large to be representative of crustal anisotropy (Barruol and Mainprice, 1993), which usually only accounts for ~ 0.1 s time delay of split teleseismic phases (e.g. Savage 1999).

3.5.3 Fossilized Lithospheric Fabrics

Because the SK(K)S raypaths sample the Nazca slab, we investigated the potential contribution of shear wave splitting from fossilized fabrics in the slab to splitting we observed at the station, which has been invoked in previous studies to explain fast polarization directions that are inconsistent with convective mantle flow models (e.g.

Conrad et al., 2007). Eakin et al. (2015) created a numerical model to simulate olivine fabrics for a segment of the subducting Nazca slab between 8°S – 20°S, just north of the Maule segment. They found that their modeled fossilized fabric orientation directions

are not consistent with observations of seismic anisotropy at the surface and concluded that fossil fabrics were overprinted during subduction and do not persist at depth. The

Nazca slab subducting in the Maule rupture zone is comparable to that modeled by

Eakin et al. (2015) in dip angle, age, and temperature, and as a result we also suggest that there is little contribution from the Nazca slab to the anisotropy observed at the

CHAMP seismic stations. Additionally, Hicks et al. (2012) investigated potential contributions from fossilized Nazca slab anisotropy to their shear wave splitting observations in the southern portion of the Maule rupture zone (~37°-39°S) and found that, with an estimated ~37 km Nazca oceanic mantle thickness, and splitting delay times greater than 1 s, the oceanic mantle would need to have a natural S-wave anisotropy in excess of 20%, a percentage that they deemed unreasonable and evidence to preclude significant contributions to teleseismic splitting from the oceanic mantle. These results, in addition to those of Eakin et al. (2015), suggest that the majority of anisotropy contributing to the shear wave splitting in the Maule segment is present either in the sub-slab mantle or the mantle wedge.

3.5.4 Trench Parallel Splitting and Sub-Slab Mantle Flow

We note a distinct lack of trench-normal splits or fast polarization directions that align with the APM of the subducting Nazca lithosphere (Figure 3-8), and therefore see little evidence for entrained supra- or sub-slab mantle flow of a-type olivine fabrics. We do not consider the model that suggests forearc b-type olivine or serpentinite fabrics and backarc a-type olivine fabrics for the mantle wedge because we do not observe a transition from trench-parallel to trench-normal splitting as we progress eastwards across the Maule segment. Instead, we suggest that the trench-parallel fast polarization directions that persist across the Maule rupture zone are consistent with the model

suggested by Russo and Silver (1994) of retrograde flow in the South American subduction zone, induced by trench migration. Kneller and van Keken (2007) modeled the effects of three-dimensional slab morphology on deformation and mantle flow in a variety of global subduction zones and found that, for Andean-type subduction systems, along-strike variations in dip angle could potentially explain trench-parallel fast polarization directions. Additionally, Lin et al. (2014) constructed dynamic subduction zone models of regional mantle flow fields for South America based on a variety of factors (e.g. slab geometry, lithosphere age, mantle composition, heterogeneities of plate thickness) and found that differential retrograde motion along-strike and influences from plate geometry can be invoked to generate trench-parallel mantle flow both in the mantle wedge and the sub-slab mantle.

3.5.5 Alignment with Nazca Slab Depth Contours

In 2014, Paczkowski et al. presented 3D kinematic-dynamic models featuring the interaction between subducting slabs and regional mantle flow and showed that deflection of regional mantle sub-slab mantle flow is a geodynamically plausible explanation for fast polarization directions observed in subduction systems with variable along-strike dip angle. In addition to the fast polarization directions determined from the

CHAMP seismic network being largely trench-parallel, we also observe splitting directions that trend sub-parallel to the contours of the subducting Nazca slab (Figure 3-

9), determined by Hayes et al. (2012) in the SLAB1.0 model based on regional seismicity, in the northern segment of the Maule rupture zone (~33°S). The gradual flattening of the Nazca lithosphere ~33°S in the northern portion of the Maule rupture zone creates accommodation space beneath the subducting lithosphere, allowing for more E-W directed mantle flow. This transition in sub-slab mantle flow regime is

manifest in our splitting observations as trench-parallel splitting south of ~33°S in regions of ~30° subduction dip, to progressively more trench-normal fast polarization directions north of ~33°S, sub-paralleling the contours of the subducting Nazca slab, and coincident with a transition in Nazca slab dip to more flat subduction. These observations suggest that the flattening of the continuous subducting Nazca slab in the northern region of the Maule segment guides the sub-slab regional mantle flow field, and are consistent with fast polarization directions oriented parallel to the contours of the subducting Nazca slab between 30°S-36°S observed by Anderson et al. (2004), who also attribute to sub-slab mantle flow guided by Nazca slab morphology.

3.6 Conclusion

We implemented the multi-channel shear wave splitting intensity method of Chevrot

(2000) to calculate the SI and shear wave splitting parameters,  and t, from 57 teleseismic events that were recorded well at the IRIS CHAMP stations. Analyses of the data set yielded measurements for 52 seismic stations from 386 event-station pairs incorporating both SKS and SKKS phases. The polarization directions we found are predominantly trench-parallel and are consistent with the retrograde flow model for subduction zone anisotropy in South America. Our teleseismic splitting measurements are not primarily generated by South American crustal anisotropy because the wavelengths of the teleseismic phases are too large to resolve crustal anisotropy and the observed time delays are much greater than expected from crustal sources.

Additionally, fossilized lithospheric fabric orientations in the subducting Nazca plate modeled around ~20°S are not consistent with regional anisotropy observed at the surface, possibly due to overprinting during subduction. Additionally, shear wave

splitting observations from previous studies in the southern portion of the Maule rupture zone (Hicks et al. 2012) are too large to be explained by anisotropic sources in the ~37 km thick Nazca oceanic mantle. We attribute the shear wave splitting we observed to anisotropy in the sub-slab mantle. Splitting fast directions observed at many CHAMP stations are rotated 40°-50° counterclockwise from the APM of the subducting Nazca lithosphere and are sub-parallel to the Nazca slab contours. This suggests that changes in the Nazca slab morphology guide the sub-slab mantle flow field.

Table 3-1. Teleseismic event locations Event Date Event Origin Latitude Longitude Depth Magnitude Magnitude Time (°) (°) (km) Type 03/20/10 14:00:49 -3.361 152.245 414.6 6.5 Mw 03/30/10 16:54:46 13.667 92.831 34.0 6.5 Mb 04/06/10 22:15:01 2.383 97.048 31.0 7.0 Mb 04/07/10 14:33:03 -3.774 141.927 33.9 6.0 Mw 04/10/10 16:54:24 -20.114 -176.223 273.2 5.9 Mw 04/11/10 09:40:25 -10.878 161.116 21.0 6.8 Mw 04/11/10 22:08:12 36.965 -3.542 6.1 6.3 Mw 04/14/10 01:25:15 33.195 96.449 7.6 6.0 Ms 04/17/10 23:15:22 -6.669 147.291 53.0 6.2 Mb 04/21/10 17:20:29 -15.271 -173.219 35.0 6.1 Mw 04/24/10 07:41:00 -1.912 128.122 27.0 6.0 Mb 04/26/10 02:59:51 22.180 123.63 14.7 6.3 Ms 05/09/10 05:59:41 3.748 96.018 38.0 6.6 Mb 05/26/10 08:53:08 25.773 129.944 10.0 6.2 Mb 05/27/10 17:14:46 -13.698 166.643 31.0 6.2 Mb 05/31/10 19:51:45 11.132 93.471 112.0 6.1 Mb 06/12/10 19:26:50 7.881 91.936 35.0 7.4 Mw 06/16/10 03:16:27 -2.174 136.543 18.0 6.7 Mb 06/16/10 03:58:08 -2.329 136.484 10.5 6.6 Mw 06/17/10 13:06:46 -33.168 179.719 170.4 6.0 Mw 06/18/10 02:23:05 44.449 148.691 25.0 5.9 Mb 06/24/10 05:32:27 -5.514 151.161 40.0 6.1 Mw 06/26/10 05:30:19 -10.627 161.447 35.0 6.6 Mw 06/30/10 04:31:02 -23.307 179.116 581.4 6.3 Mw 07/02/10 06:04:03 -13.643 166.485 29.0 6.3 Mw 07/04/10 21:55:51 39.697 142.369 27.0 6.4 Mb 07/10/10 11:43:32 11.143 145.999 13.0 6.2 Mb 07/18/10 13:04:09 -5.966 150.428 28.0 6.3 Mb 07/20/10 19:18:21 -5.914 150.702 35.0 6.3 Mw 07/21/10 09:16:04 3.039 128.222 100.0 6.0 Mb 07/22/10 05:03:57 -15.149 168.168 10.0 5.9 Mw 07/23/10 22:08:11 6.486 123.467 585.8 6.9 Mb 07/23/10 22:08:11 6.718 123.409 607.1 6.3 Mb 07/24/10 05:35:01 6.218 123.519 553.0 6.6 Mw 07/25/10 03:39:17 -15.069 -173.548 7.0 5.9 Mb 07/30/10 03:56:13 52.498 159.843 23.0 6.3 Mw 08/03/10 12:08:25 1.239 126.213 41.0 5.9 Ms 08/04/10 12:58:24 51.423 -178.649 27.0 6.2 Mb 08/04/10 04:46:20 -26.916 -177.244 18.0 5.9 Mw 08/04/10 07:15:34 -5.496 146.811 225.6 6.1 Mb 08/04/10 22:01:43 -5.746 150.765 44.0 6.9 Mw 08/10/10 05:23:44 -17.541 168.069 25.0 7.3 Mw 08/10/10 23:18:31 -14.460 167.345 191.6 5.9 Mw 08/13/10 21:19:33 12.484 141.476 10.0 6.5 Mb

Table 3-1. Continued Event Date Event Origin Latitude Longitude Depth Magnitude Magnitude Time (°) (°) (km) Type 08/14/10 23:01:04 12.273 141.429 13.0 6.3 Ms 08/14/10 07:30:16 12.348 141.487 10.0 6.2 Mw 08/15/10 15:09:29 -5.692 148.342 174.7 6.3 Mw 08/16/10 19:35:49 -20.799 -178.826 603.0 6.2 Mw 08/16/10 03:30:53 -17.759 65.647 9.8 6.3 Mw 08/18/10 16:28:15 12.234 141.514 10.0 6.3 Mw 09/03/10 11:16:06 51.450 -175.87 23.5 6.3 Mw 09/03/10 16:35:47 -43.520 171.83 12.0 7.0 Mw 09/04/10 08:52:04 -17.367 -173.998 69.0 5.9 Mb 09/07/10 16:13:32 -15.880 -179.304 10.0 6.0 Ms 09/08/10 11:37:32 -20.664 169.803 10.0 6.2 Ms 09/17/10 19:21:15 36.440 70.770 220.1 6.3 Mw 09/26/10 12:12:41 -5.312 133.928 30.0 6.1 Mb

Table 3-2. Teleseismic splitting measurements Latitude Longitude Phi Dt No. Station RMS (°) (°) (°) (s) Evts u01b -37.29 -72.49 19.40  2.23 1.11  0.08 2.09 12 u02b -37.21 -72.98 35.80  2.97 1.44  0.18 2.48 4 u03b -37.70 -72.33 18.20  2.20 1.52  0.08 1.24 6 u04b -37.99 -72.57 45.70  0.96 1.83  0.12 1.73 13 u05b -37.95 -72.81 42.90  3.23 0.91  0.08 1.81 12 u06b -37.22 -73.55 19.30  0.79 1.91  0.06 0.71 11 u07b -38.25 -73.47 2.40  1.51 2.39  0.16 2.02 13 u08b -36.63 -72.59 56.30  1.23 1.46  0.08 1.24 12 u09b -38.49 -73.18 14.30  2.56 1.42  0.13 0.90 9 u10b -38.20 -72.85 17.00  1.16 0.87  0.04 1.62 13 u11b -37.21 -71.83 18.00  1.72 1.55  0.10 0.95 5 u12b -37.95 -73.41 31.90  3.04 2.42  0.09 1.34 6 u15b -38.07 -73.00 32.10  1.80 0.77  0.07 1.63 6 u16b -37.82 -72.96 62.30  4.18 1.66  0.14 1.90 7 u26b -36.52 -72.22 38.70  0.35 2.02  0.06 1.53 8 u27b -36.28 -72.53 42.30  1.70 1.76  0.25 1.63 2 u28b -36.33 -72.33 58.50  0.75 1.74  0.07 1.53 4 u29b -36.91 -71.50 9.40  1.76 1.31  0.06 0.98 7 u30b -36.80 -71.76 16.10  9.46 1.89  0.80 2.45 3 u32b -36.33 -71.74 40.30  0.38 2.44  0.08 0.89 5 u33b -36.56 -71.54 8.90  0.98 1.61  0.05 1.08 9 u34b -36.45 -71.71 4.20  1.16 1.79  0.06 0.38 7 u35b -36.90 -71.85 45.80  6.15 3.14  0.63 0.43 4 u36b -36.76 -72.37 37.90  1.07 1.51  0.13 1.48 9 u37b -36.80 -72.62 -50.1  0.48 1.09  0.05 1.61 6 u40b -36.63 -72.86 9.80  1.41 1.42  0.06 0.41 5 u41b -36.40 -72.84 10.10  3.46 1.19  0.15 1.05 2

Table 3-2. Continued Latitude Longitude Phi Dt No. Station RMS (°) (°) Evts u42b -36.91 -72.84 39.60  0.42 2.10  0.06 1.24 7 u43b -36.22 -71.48 9.00  1.17 1.85  0.05 1.02 10 u44b -36.09 -72.11 76.60  1.92 0.62  0.02 1.37 7 u45b -34.74 -71.64 -19.9  0.88 3.45  0.11 0.77 4 u46b -35.06 -71.15 23.10  2.87 1.18  0.10 2.42 6 u51b -34.65 -71.21 40.50  0.64 1.67  0.06 1.38 9 u52b -35.80 -71.44 33.20  2.73 1.26  0.08 2.98 7 u54b -35.55 -71.36 -10.6  0.87 2.05  0.09 0.52 4 u55b -35.70 -71.10 44.40  0.61 1.59  0.08 3.42 4 u56b -35.29 -71.25 -7.00  1.30 2.00  0.06 0.79 5 u57b -35.49 -72.10 43.00  1.52 1.11  0.08 0.80 8 u58b -35.37 -72.45 34.00  2.22 1.48  0.16 1.73 7 u59b -35.91 -72.40 31.60  1.26 1.25  0.11 1.91 7 u60b -35.88 -72.62 41.90  0.55 1.46  0.09 1.20 10 u61b -35.57 -72.60 -33.6  0.84 1.57  0.07 1.75 8 u62b -35.66 -71.77 48.10  0.96 3.46  0.09 1.82 4 u63b -35.73 -72.13 51.10  0.38 2.85  0.08 1.38 9 u64b -35.40 -71.68 -0.38  2.26 1.72  0.08 0.70 2 u66b -34.91 -72.17 8.12  1.92 1.09  0.04 1.93 10 u67b -34.49 -71.55 28.60  0.94 1.05  0.04 1.30 12 u68b -34.41 -71.97 -6.27  0.97 1.49  0.08 0.69 13 u70b -33.77 -70.87 39.50  0.44 2.01  0.05 0.91 11 u72b -33.98 -71.81 19.70  4.74 0.88  0.11 2.34 7 u74b -33.60 -71.54 14.60  13.20 0.31  0.12 1.85 6 u75b -33.38 -71.18 43.50 1.98 1.23  0.10 1.68 9

Figure 3-1. Schematic representation of Russo and Silver (1994) of retrograde flow model for the Nazca-South America subduction zone.

Figure 3-2. Teleseismic event locations. The black star represents the Maule, Chile study area and concentric circles around the star denote epicentral distances from the study area in increments of 30°.

Figure 3-3. Radial and transverse component records to illustrate Chevrot (2000) shear wave splitting intensity method. Recording from CHAMP station U51B of a magnitude 6.3 earthquake with origin time of July 20, 2010 at 19:39:13.640. Event latitude is -5.914° and event longitude is 150.702°, with a depth of 35 km. Red lines indicate the window used for the multichannel approach. Pink dashed lines represent automated phase arrival picks determined from a radial earth velocity model.

Figure 3-4. Zooming in on the window selected in Figure 3-3, the dashed red line represents the derivative of the radial component multiplied by a coefficient of splitting and projected onto the transverse component, shown in blue.

Figure 3-5. Backazimuth vs SI plot. Each red dot represents an event-station pair coefficient of splitting measurement at the respective azimuth. The blue line is the best-fit sinusoidal curve for these data.

Figure 3-6. Rose histogram of the teleseismic fast directions for the entire dataset. Black arrow represents mean direction of dataset.

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Figure 3-7. Teleseismic shear wave splitting measurements from SK(K)S phases recorded at stations in the IRIS CHAMP temporary seismic network. Red triangles represent seismic stations. White arrows represent the Nazca APM. Red arrows denote trench-normal directions for portions of the Maule segment. Blue splitting vectors are rotated clockwise from Nazca APM, yellow splitting vectors are rotated counterclockwise from Nazca APM, and purple splitting vectors are teleseismic measurements from CHARGE stations (Anderson et al., 2004).

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Figure 3-8. Histogram of splitting directions. Yellow bars are directions counterclockwise from Nazca APM and blue bars are directions clockwise from Nazca APM. Dominant splitting direction is ~40°-50° counterclockwise from Nazca APM. There are no fast directions that are present in the trench normal direction for the Maule segment.

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Figure 3-9. Shear wave splitting measurements plotted with depth contours to the top of the subducting Nazca lithosphere from the SLAB1.0 model (Hayes et al., 2012).

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CHAPTER 4 SHEAR WAVE SPLITTING OF MAULE AFTERSHOCKS

4.1 Background

It is not trivial to isolate the contributions to observed shear wave splitting that accrue from propagation through the sub-slab mantle, the Nazca slab, the suprasubduction mantle wedge, and the overriding South American crust, each of which may generate distinct splitting of the teleseismic shear waves used in the study presented in Chapter 3. To discriminate between anisotropy generated in two of three distinct structural volumes, the mantle wedge and overriding crust, we compiled a dataset of aftershocks from the Maule event that nucleated along the slab interface or in the overriding South American crust. We then made shear wave splitting intensity measurements using the local S phases from these events recorded at IRIS CHAMP stations. The upgoing local/regional S phase has much shorter travel paths – restricted to the mantle wedge and/or South American crust – and also much shorter wavelengths

(~2-3 km) than the teleseismic SK(K)S phases used in Chapter 3, making them sensitive to seismic anisotropic fabrics present only in these shallower portions of the subduction zone, and also rendering these waves sensitive to much smaller structures and finer anisotropic fabrics. The study described here incorporates more data than could be used as input to the standard shear wave splitting techniques, as applied and described in previous studies in the region (MacDougall et al., 2012), and we show that the SI method also yields many more robust measurements.

4.2 Data and Methods

We selected 132 local events (Rietbrock et al., 2012) that were recorded at the

IRIS CHAMP broadband seismic stations (Table 4-1), all with magnitudes greater than

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3.5, to assure a high signal to noise ratio. Because we are using upgoing S waves generated by local seismic events, we cannot assume that the fast wave is radially polarized, as was the case for the SK(K)S phases, and instead we use the S wave polarization direction observed at the station as the assumed initial polarization direction. To assure that there is no interference with crustal reverberations and to avoid potential modification of the phase of the S arrivals due to shallow angle incidence at the Earth’s surface, we limited the incidence angle of the upgoing S waves to be no greater than 75° from vertical within the upper 6 km of the crust, an extension of the allowable shear wave window, as defined by Booth and Crampin (1985), which we invoke following Hung and Forsyth (1999); all suitable waveforms were then filtered between 0.3 and 3.5 Hz. This dataset samples a much shallower volume than the data we used for the SK(K)S study, given that the hypocenter of the deepest event we used was only 66 km deep, and the longest event-station path was 273 km. These selection criteria yielded 233 event-station pair measurements recorded at 39 stations after we applied the shear wave splitting intensity method of Chevrot (2000) modified for a local dataset, described in more detail in Chapter 3.3. With teleseismic events, Chevrot

(2000) projected the transverse components on the radial components derivative, assumed that the incoming shear wave was radially polarized, and determined the SI as a function of backazimuth. For local events, we cannot make the assumption that the incoming shear wave is radially polarized prior to incidence on an anisotropic region, and instead must determine the incoming polarization angle. We applied the Silver and

Chan (1991) method to calculate the polarization angle observed at the station and use that angle as the assumed polarization angle prior to incidence on an anisotropic region.

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When extracting the splitting parameters, we plot the SI as a function of incoming polarization angle, unlike the core-refracted teleseismic events used in Chapter 3 which are plotted as a function of backazimuth.

4.3 Results

The splitting directions determined in this study reveal a complicated geographic pattern of fast directions. Despite the apparent scatter in the dataset, fast directions trend in three dominant directions, N40°E, N20°W, and N10°W-N20°E (Figure 4-2 and

Figure 4-3). We observe a subset of fast polarization directions trending ~N60°-90°E.

Stations U65B and U68B had the most measurements (each with 14 useable waveforms), both of which are located in the Coastal Cordillera. Most IRIS CHAMP stations located in the Central Depression did not record enough suitable waveforms to produce robust splitting measurements. Delay times range from 0.025 s to 0.358 s and show no spatial pattern within the rupture zone. The errors associated with the delay times, however, are often more than 100% of the dt value and make interpretation of the splitting times in a geologic context difficult.

4.4 Discussion

The wavelengths of the local S waves within our restricted frequency band (0.3-

3.5 Hz) are ~2-3 km and potentially allow us to resolve features of a similar scale in the mantle wedge and overriding South American plate. The three dominant fast polarization directions we obtained in our dataset are N40°E, N20°W, and N10°W-20°E, with a subset of measurements trending N60°-90°E. Using the multichannel splitting intensity approach of Chevrot (2000), it is not possible to assess the fast polarization direction and time delay for each event-station pair, as the splitting parameters at each station are determined based on fitting a curve to many SI/initial polarization angle pairs

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recorded at each individual station. We did, however, consider the average path-length recorded at each station and compared those with the time delay to see if there was any relationship between these two parameters, as would be expected if delay time accrued systematically due to propagation through a consistently oriented anisotropic fabric.

However, we do not observe a robust relationship between average path-length and the splitting delay at any of the stations in the IRIS CHAMP network for the local events we used.

4.4.1 Supra-slab Entrained Mantle Flow

With hypocenters extending to depths of 66 km (epicenter at -33.438°S, -

71.050°W, nucleating in the mantle wedge) in this dataset of Maule aftershocks, one potential explanation for splitting measurements trending ~N60°-90°E (plotted as red bars in Figure 4-3) is that they are sampling anisotropy located in the supra-slab mantle, which is entrained in corner wedge flow in the subduction zone. This follows a 2D

‘corner flow’ model (e.g., Ida 1983; Hall et al., 2000) with flow lines approximately paralleling the APM of the subducting Nazca oceanic plate.

4.4.2 Contributions from Crustal Anisotropy

Previous studies of local shear wave splitting measurements have attributed observed shear wave splitting to crustal anisotropy primarily on the basis of small delay times indicative of travel through a (relatively) thin crust and fast polarization directions that parallel ambient tectonic stress directions or the strikes of active faults (e.g.,

Christensen, 1966a,b; Nur and Simmons, 1969, Crampin 1977, 1981; Kaneshima 1990;

Helffrich et al., 1994; McNamara et al., 1994; Chang et al., 2009; Huang et al., 2011).

Some studies suggest that seismic anisotropy in the crust can be attributed to rock cracks and fractures (e.g. Hudson 1981, 1994; Schoenberg and Douma 1988). Global

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reviews of crustal anisotropy (e.g. Crampin 1994) found that, while the upper 10-15 km of the crust can contain anisotropic fabrics due to cracks and microcracks (Kaneshima et al., 1988; Savage 1999), the majority of fractured anisotropic rocks are concentrated in the uppermost 5 km of the crust (Gaiser and Carrigan 1990; Gledhill 1990;

Kaneshima 1990). Cracks in the lower crust and the uppermost mantle do not produce observed anisotropy because cracks at these depths are closed due to increasing overburden (Levin and Park 1997), as shown in laboratory experiments in which, when exposed to pressures greater than 200-300 MPa corresponding to depths greater than

15 km, cracks close (e.g.. Kern, 1990; Hrouda et al., 1993). Others have proposed that observed splitting fast polarization directions may parallel layer foliation or lithological lineations (e.g., Okaya et al., 1995; Godfrey et al., 2000; Meltzer and Christensen

2001). Some authors (e.g., Barruol and Kern 1996; Godfrey et al., 2000) suggest that ductile flow in the lower crust could align minerals, such as mica, parallel to the slow symmetry axis of the anisotropic material, resulting in fast directions that are orthogonal to the dominant shear direction. Analyses of crustal anisotropy from shear wave splitting can be useful in determining the state of stress of the upper crust.

We expect crustal petrofabrics to strongly influence our local anisotropy measurements because the majority of the earthquakes in this study nucleated within the South American continental crust (Table 4-1) and have ray paths that traverse the

South American continental crust almost exclusively. The crustal thickness of the

Coastal Cordillera in the Maule segment increases southwards, from 30-45 km thick

(Alvarado et al., 2007) in the north, to 45-50 km thickness (Hicks et al., 2014) in the south. Estimates from receiver functions, however, place the depth to the base of the

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continental Moho shallower at ~36.5°S, extending to ~38 km depth (Dannowski et al.,

2013). Progressing eastwards to the Central Depression, crustal thickness estimates are ~30 km thick (Krawczyk et al., 2006), with thickness increasing towards the Main

Andean Cordillera to ~40-65 km (Allmendinger et al., 1983; Krawczyk et al., 2006).

Pervasive, regional crustal foliation directions potentially present in the Central

Depression crust are hidden beneath Quaternary sediments. The Main Andean

Cordillera in the Maule rupture zone is characterized by weakly deformed fabrics. The

Eastern and Western Metamorphic Series of the Chilean Coastal Cordillera, however, do have distinct foliation patterns that are readily observable regionally from distributed outcrops. This is complicated, however, by the variable convergence directions and rates throughout the evolution of the Chilean subduction system (e.g., Pardo-Casas and

Molnar 1987; Yanez et al., 2002) and complicates correlations between crustal fabric direction and the history of plate convergence direction.

Two of the dominant SI directions we observe, N40°E and N20°W, are consistent with pervasive development of conjugate shears in crustal or shallow upper mantle rocks sampled by the upgoing S waves from the Maule aftershocks. These observed SI fast directions (Figure 4-3), are 60° from one another, the classic angle spread between conjugate shears developed in pure shear. Thus, we consider these two SI observation groups to be consistent with the development of near vertical flattening fabric due to long term convergence and compression during subduction of the Nazca (or previously subducting Farallon) plate. These conjugate shearing directions also imply along-strike compression, perhaps as a result of oroclinal bending (Russo and Silver, 1996).

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We lack robust splitting measurements at seismic stations located in the Central

Depression (Figure 4-1). These stations were sited in the sedimentary basin and, as a result, are not coupled well with the underlying basement rock. Consequently, it is not surprising that the seismic stations located on the 1-2 km thick Quaternary sedimentary fill of this Longitudinal Valley do not yield robust shear wave splitting measurements, as we expect many of our local fast polarization directions are sourced from crustal anisotropy.

4.4.3 Lateral Mantle Flow in the Forearc

Fast polarization directions oriented between N10°W-N20°E are approximately parallel to the strike of the subduction trench in the Maule segment. We interpret these fast directions to be representative of supra-slab mantle flow that is migrating laterally along strike, around the subducting Nazca slab beneath the forearc. Farther to the west, the flow is more consistent with the 2D corner flow models in which entrained mantle flow is induced by down dip motion of the slab, as described in Chapter 4.4.1. Our interpretation is consistent with the model proposed by Long and Silver (2008) (Figure

4-4), which they developed after observations in 13 global subduction zones. Long and

Silver (2008) suggested that mantle wedge flow is a combination of the classic corner flow induced by convergence and lateral flow along-strike near the forearc induced by trench motion, characterized by trench-parallel splitting close to the trench with a transition to trench-perpendicular fast polarization directions, or directions paralleling the

APM of the subducting plate, in the backarc of the subduction system.

4.4.4 Robustness of Multichannel Analysis Method for Local Datasets

Our measurements, using the Chevrot (2000) method, are more consistent than those obtained from the shear wave splitting study of MacDougall et al. (2012), which

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also incorporated local S phases recorded at stations in the IMAD network in the Maule rupture zone, but instead implemented the eigenvalue minimization method of Silver and Chan (1991). The highly variable fast polarization directions that MacDougall et al.

(2012) found were very scattered, making them difficult to interpret (Figure 4-5). A comparison of our results (Figure 4-1) and those of MacDougall et al. (2012) (Figure 4-

5) clearly shows that the multichannel technique yields a larger, more consistent data set when analyzing local waveforms for shear wave splitting.

4.4.5 Comparison to Teleseismic Dataset

Unlike the SK(K)S results, the upgoing S splitting intensity measurements are characterized by much greater variability. We see delay times that are much shorter than for the teleseismic phases which could be a result of the different frequency bands used for filtering, but is more likely due to the fact that the datasets sample different volumes along different path lengths (i.e., the teleseismic waves sample the sub-slab, slab, mantle wedge and overlying continental crust whereas the local waves sample the mantle wedge and overlying continental crust exclusively). The greater splitting delay time observed for the teleseismic events suggest that the majority of the anisotropy the teleseismic phases encounter is located below the local event hypocenters. This is consistent with our interpretation that the anisotropy observed in the teleseismic dataset is likely sourced subslab, with fast polarization directions that are predominantly trench- parallel due to the retrograde trench motion. The fast polarization directions from the local events serve as an ancillary dataset to the teleseismic results and, when used together, can help better constrain regions of anisotropy in the subduction zone.

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4.6 Conclusion

We implemented the multi-channel shear wave splitting intensity method of

Chevrot (2000) to calculate the shear wave splitting parameters  and t from 132 aftershocks from the Maule 2010 earthquake. Analyses of the dataset yielded measurements for 39 seismic stations using 233 event-station pairs incorporating the local S phase, with stations U65B and U68B having the most measurements (with 14 useable waveforms), both of which are located in the Coastal Cordillera. We observe three dominant splitting directions with fast polarization orientations of N40°E, N20°W, and N10°W-N20°E and a subset of splitting directions between N60°-90°E. This distribution of fast polarization directions is consistent with shear deformation in the direction in maximum compressive stress (N80°E) due to Late Cenozoic subduction of the Nazca plate, and the development of flattening and shear as a result of long-term conjugate shearing of crustal South American rocks due to subduction. This suggests that the splitting observed in our measurements are sourced either in the supra-slab mantle flow or crust for the N60°-90°E fast direction subset, and are likely of crustal origin for the N40°W and N20°E subsets of fast polarization directions. The paucity of measurements in the Central Depression supports the argument of crustally sourced anisotropy because these seismic stations are seated on 1-2 km of sedimentary basin infill. Additionally, fast polarization directions oriented N10°W-N20E suggest potential lateral mantle flow along-strike, adhering to the mantle wedge flow model proposed by

Long and Silver (2008) in which mantle flow is a combination of 2D corner flow and along-strike lateral flow beneath the forearc.

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Table 4-1. Local event locations Event Event Origin Latitude Longitude Depth Magnitude Magnitude Date Time (°) (°) (km) Type 03/29/10 08:19:06 -36.7619 -73.0523 27.7 4.4 Mw 04/01/10 12:53:07 -34.6988 -72.0676 9.88 5.6 Mw 04/01/10 19:34:09 -34.9028 -71.8718 46.4 4.5 Mw 04/01/10 19:58:58 -34.1045 -72.3744 23.5 4.2 Mw 04/02/10 10:38:22 -36.6527 -73.5333 19.29 5.4 Mw 04/02/10 14:58:10 -34.7765 -71.6748 46.83 5.1 Mw 04/02/10 19:34:09 -36.1243 -72.8662 30.56 5.5 Mw 04/02/10 22:58:07 -36.2621 -73.1193 26.7 6.2 Mw 04/03/10 02:12:39 -36.2822 -73.1461 29.3 5.1 Mw 04/05/10 03:32:15 -33.4382 -71.0495 66.0 5.4 Mw 04/05/10 13:21:06 -37.4025 -72.8405 44.3 4.5 Mw 04/06/10 05:38:38 -38.1577 -73.3197 28.5 4.2 Mw 04/06/10 13:04:40 -35.0574 -72.9386 24.0 5.7 Mw 04/07/10 02:18:36 -35.5530 -72.8310 22.6 5.0 Mw 04/07/10 02:29:48 -35.5966 -72.7918 9.3 5.3 Mw 04/07/10 16:25:33 -34.7402 -71.8071 48.2 5.4 Mw 04/07/10 17:50:40 -37.0626 -73.2575 25.3 4.3 Mw 04/07/10 18:34:21 -34.8060 -71.7811 22.6 5.0 Mw 04/08/10 21:15:04 -34.4168 -72.0967 17.2 5.3 Mw 04/08/10 22:21:12 -36.6189 -73.7487 23.7 5.3 Mw 04/09/10 00:30:29 -38.2279 -73.1813 46.1 5.3 Mw 04/11/10 02:34:52 -35.9759 -72.7285 30.3 4.2 Mw 04/11/10 02:43:45 -35.6199 -72.7095 23.8 4.4 Mw 04/11/10 10:28:07 -34.8047 -71.8008 49.0 5.5 Mw 04/12/10 09:57:40 -35.5571 -72.3661 29.4 5.1 Mw 04/12/10 22:09:59 -34.9293 -71.5918 17.7 3.6 Mw 04/13/10 15:10:38 -33.9835 -72.2504 20.8 5.4 Mw 04/13/10 15:13:07 -33.9606 -72.2364 21.9 5.2 Mw 04/14/10 04:46:54 -35.3104 -73.2048 20.6 5.2 Mw 04/15/10 08:20:21 -35.1487 -71.7585 52.3 5.1 Mw 04/15/10 10:13:11 -34.7039 -72.0370 9.6 5.0 Mw 04/15/10 12:12:46 -36.5566 -72.8676 32.1 5.2 Mw 04/16/10 23:15:34 -37.4727 -74.1044 29.3 5.5 Mw 04/18/10 02:52:16 -37.2261 -74.1546 28.4 5.0 Mw 04/18/10 10:01:59 -35.9782 -72.7616 27.4 4.2 Mw 04/19/10 07:21:31 -35.7198 -72.4661 28.6 5.1 Mw 04/19/10 07:32:45 -37.6011 -74.1528 36.9 5.6 Mw 04/20/10 07:41:18 -33.9794 -71.4200 45.2 4.4 Mw 04/23/10 05:09:04 -35.7673 -72.8519 27 4.0 Mw 04/23/10 10:03:05 -37.6340 -73.2223 43.7 6.0 Mw 04/25/10 07:06:54 -36.3902 -72.4155 45.1 4.2 Mw 04/25/10 10:12:48 -34.0317 -72.2424 18.2 5.1 Mw 04/25/10 15:42:22 -37.6461 -73.1656 34.0 4.7 Mw 04/25/10 17:29:14 -36.2071 -72.9977 24 4.1 Mw 04/26/10 08:44:29 -37.4874 -73.0286 31.4 4.6 Mw

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Table 4-1. Continued Event Event Origin Latitude Longitude Depth Magnitude Magnitude Date Time (°) (°) (km) Type 04/27/10 05:37:18 -34.3737 -72.0549 13.6 5.4 Mw 04/28/10 04:26:30 -35.7232 -72.8318 30.1 5.1 Mw 04/29/10 07:21:02 -37.8959 -73.0796 36.9 4.1 Mw 04/29/10 08:38:51 -34.9848 -72.6619 25.8 5.0 Mw 04/29/10 08:39:25 -35.1524 -71.8782 47.1 5.0 Mw 04/29/10 13:40:12 -36.6043 -73.3222 26.1 5.3 Mw 04/29/10 16:38:53 -34.0916 -71.9977 25.0 4.0 Mw 04/30/10 09:16:39 -34.2769 -72.2658 23.5 4.2 Mw 05/01/10 14:41:06 -33.2784 -72.3230 25.0 5.4 Mw 05/02/10 14:52:41 -34.4439 -72.2619 11.5 6.1 Mw 05/02/10 19:22:58 -34.2788 -72.1717 17.6 5.1 Mw 05/02/10 19:46:29 -34.3177 -72.1706 22.6 5.2 Mw 05/03/10 06:33:28 -34.2732 -72.1739 4.2 5.2 Mw 05/03/10 06:49:26 -34.2898 -72.1737 24.2 5.4 Mw 05/03/10 10:22:27 -37.0814 -72.7454 33 3.7 Mw 05/03/10 18:39:40 -37.2480 -74.1034 29.1 5.2 Mw 05/03/10 23:09:46 -37.9870 -73.7093 35.2 5.9 Mw 05/04/10 13:55:06 -34.1589 -72.4609 16.7 5.5 Mw 05/05/10 03:32:35 -33.4747 -72.3234 31.2 5.2 Mw 05/05/10 10:29:30 -35.4542 -72.6113 28.4 5.3 Mw 05/05/10 15:24:07 -35.5980 -73.3984 17.7 5.9 Mw 05/06/10 05:43:22 -35.1968 -72.4220 28.6 5.1 Mw 05/08/10 01:43:20 -34.3407 -72.4362 18.0 5.0 Mw 05/09/10 03:29:36 -34.0114 -72.3593 18.8 5.7 Mw 05/09/10 10:44:56 -37.5629 -73.9634 29.9 5.3 Mw 05/09/10 23:05:26 -34.0464 -72.3186 26.4 5.1 Mw 05/11/10 13:04:53 -35.6563 -73.1019 12.0 5.3 Mw 05/13/10 03:56:03 -35.1323 -73.1076 17.9 5.4 Mw 05/13/10 16:32:48 -35.0053 -72.0714 44.6 4.1 Mw 05/13/10 20:39:12 -34.2581 -72.4473 23.5 5.5 Mw 05/13/10 22:56:59 -37.0489 -73.3732 26.2 4.4 Mw 05/15/10 04:14:19 -34.9447 -71.6500 49.1 4.0 Mw 05/15/10 08:06:32 -38.1722 -73.7450 27.6 4.6 Mw 05/17/10 00:22:22 -37.8091 -73.2180 31.6 4.1 Mw 05/17/10 18:24:57 -37.1633 -74.0955 26.7 5.3 Mw 05/17/10 21:16:42 -36.6667 -73.1385 30.7 5.3 Mw 05/18/10 10:11:20 -34.8679 -72.9431 14.7 5.4 Mw 05/19/10 06:01:02 -34.3437 -71.9923 22.3 3.7 Mw 05/19/10 09:38:48 -34.2944 -72.3959 28.9 5.4 Mw 05/19/10 10:48:05 -34.6796 -72.5815 17.9 5.3 Mw 05/19/10 13:45:32 -36.7318 -73.8563 12.7 5.3 Mw 05/20/10 16:53:40 -35.6274 -73.0468 22.8 5.0 Mw 05/21/10 18:52:08 -34.6347 -71.9943 11.9 5.9 Mw 05/23/10 00:24:07 -37.5511 -73.7839 24.4 4.1 Mw 05/24/10 14:18:06 -34.4050 -71.7611 46.4 5.3 Mw

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Table 4-1. Continued Event Event Origin Latitude Longitude Depth Magnitude Magnitude Date Time (°) (°) (km) Type 05/24/10 19:12:14 -35.7247 -73.0234 9.5 5.6 Mw 05/25/10 13:09:25 -37.7162 -73.2289 31.6 5.2 Mw 05/25/10 15:56:06 -35.5539 -72.5552 29.0 3.8 Mw 05/26/10 07:10:57 -33.7492 -72.0583 23.8 4.0 Mw 05/27/10 03:54:02 -34.7645 -71.6907 46.4 4.4 Mw 05/28/10 19:23:13 -34.7356 -71.8107 41.2 4.2 Mw 06/08/10 14:33:03 -34.7293 -71.8431 45.8 4.3 Mw 06/11/10 00:52:38 -34.7578 -71.7220 49.2 4.4 Mw 06/19/10 01:08:36 -33.9145 -71.5456 40.3 3.7 Mw 06/20/10 14:21:35 -37.6229 -73.7358 24.3 3.7 Mw 06/29/10 01:40:01 -37.8467 -73.6838 24.5 5.4 Mw 07/01/10 20:58:23 -35.6688 -72.5989 29.3 4.8 Mw 07/03/10 00:05:07 -34.7508 -71.7553 46.2 5.0 Mw 07/06/10 13:54:02 -35.6730 -72.1085 45.4 4.8 Mw 07/08/10 02:11:36 -34.8022 -71.7269 38.0 4.4 Mw 07/14/10 14:21:20 -35.0985 -72.4483 24.2 4.1 Mw 07/16/10 12:47:37 -37.7382 -73.2362 32.0 4.3 Mw 07/16/10 19:30:44 -38.2168 -73.7281 17.2 4.2 Mw 07/17/10 01:39:40 -35.5044 -72.8447 21.6 4.1 Mw 07/18/10 01:12:25 -35.1406 -71.8655 48.9 4.1 Mw 07/20/10 19:05:02 -38.0154 -73.6013 21.8 4.1 Mw 07/22/10 11:27:49 -37.2368 -72.7416 40.6 4.3 Mw 07/23/10 13:50:55 -37.0775 -73.6460 18.2 4.3 Mw 07/30/10 12:10:54 -37.3476 -73.7911 18.7 4.3 Mw 08/06/10 15:15:51 -35.0105 -72.6161 26.0 4.5 Mw 08/06/10 21:05:27 -38.2391 -73.8352 16.4 4.6 Mw 08/09/10 12:13:43 -38.5508 -73.2206 34.6 4.8 Mw 08/09/10 15:51:03 -37.7378 -73.3014 22.7 4.3 Mw 08/14/10 14:40:22 -34.7383 -71.8128 43.8 4.2 Mw 08/25/10 13:14:38 -35.0766 -72.2225 26.2 4.3 Mw 08/25/10 22:34:54 -37.9272 -73.5565 23.4 4.2 Mw 08/28/10 04:48:03 -34.7723 -71.8472 41.0 4.2 Mw 08/30/10 08:34:48 -35.2504 -72.0838 17.8 4.6 Mw 09/01/10 17:08:53 -36.8238 -73.1269 24.9 4.2 Mw 09/08/10 01:30:20 -34.7731 -71.8505 44.3 4.4 Mw 09/12/10 19:16:22 -33.9565 -71.4417 41.9 4.6 Mw 09/13/10 05:16:59 -34.7418 -71.8138 43.6 4.2 Mw 09/14/10 05:43:09 -34.1930 -72.2248 26.9 3.8 Mw 09/15/10 05:03:34 -34.8910 -72.3015 28.8 4.1 Mw 09/15/10 08:52:25 -33.6604 -71.9189 24.9 4.1 Mw 09/17/10 05:52:02 -33.6163 -72.0637 23.0 4.3 Mw 09/23/10 16:36:17 -34.9773 -71.8394 44.3 5.3 Mw

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Table 4-2. Local splitting measurements Latitude Longitude No. Station (°) (°) Phi Dt RMS Evts u02b -37.21 -72.98 -36.55  0.31 0.140  0.117 0.145 6 u06b -37.22 -73.55 -35.49  0.61 0.232  0.635 0.111 3 u07b -38.25 -73.47 60.15  0.70 0.109  0.142 0.213 7 u08b -36.63 -72.59 79.31  0.44 0.358  0.808 0.035 3 u10b -38.20 -72.85 45.09  0.34 0.138  0.423 0.009 3 u12b -37.95 -73.41 -9.66  0.45 0.095  0.251 0.023 3 u15b -38.07 -73.00 12.57  0.75 0.146  0.376 0.104 4 u16b -37.82 -72.96 40.66  0.15 0.089  0.057 0.024 4 u28b -36.33 -72.33 -25.16  0.45 0.025  0.053 0.056 7 u29b -36.91 -71.50 35.41  0.44 0.132  0.233 0.061 3 u30b -36.80 -71.76 26.32  0.84 0.043  0.078 0.047 3 u35b -36.90 -71.85 84.64  0.12 0.162  0.291 0.228 3 u37b -36.80 -72.62 -10.16  0.82 0.040  0.122 0.045 5 u40b -36.63 -72.86 -84.63  2.19 0.026  0.144 0.099 6 u41b -36.40 -72.84 -51.28  0.17 0.115  0.126 0.106 6 u42b -36.91 -72.84 65.12  0.52 0.051  0.161 0.051 4 u43b -36.22 -71.48 -89.07  0.42 0.066  0.103 0.118 7 u44b -36.09 -72.11 -5.29  0.18 0.114  0.222 0.058 3 u45b -34.74 -71.64 74.60  0.74 0.033  0.096 0.063 9 u46b -35.06 -71.15 1.93  1.44 0.026  0.093 0.067 8 u51b -34.65 -71.21 -2.01  0.23 0.133  0.112 0.071 6 u54b -35.55 -71.36 -37.77  1.40 0.026  0.073 0.087 8 u56b -35.29 -71.25 33.72  0.22 0.170  0.214 0.004 3 u57b -35.49 -72.10 8.22  1.36 0.032  0.140 0.032 8 u58b -35.37 -72.45 80.39  0.76 0.017  0.069 0.090 9 u59b -35.91 -72.40 39.40  0.26 0.174  0.248 0.081 5 u60b -35.88 -72.62 -28.41  0.56 0.042  0.113 0.061 6 u61b -35.57 -72.60 -52.44  1.51 0.054  0.210 0.015 4 u62b -35.66 -71.77 48.57  0.24 0.123  0.213 0.052 5 u63b -35.73 -72.13 -16.35  0.37 0.072  0.139 0.072 6 u64b -35.40 -71.68 53.29  0.36 0.039  0.073 0.063 8 u65b -34.96 -71.79 29.87  0.39 0.033  0.029 0.092 14 u66b -34.91 -72.17 4.57  0.70 0.081  0.200 0.045 3 u67b -34.49 -71.55 16.00  0.26 0.123  0.145 0.064 6 u68b -34.41 -71.97 70.12  0.22 0.061  0.076 0.061 14 u69b -34.37 -71.18 -73.25  0.22 0.062  0.061 0.088 11 u72b -33.98 -71.81 -28.12  0.49 0.083  0.127 0.069 6 u73b -33.92 -71.42 9.55  0.41 0.042  0.053 0.034 8 u74b -33.60 -71.54 30.53  0.60 0.029  0.049 0.064 6

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Figure 4-1. Local splitting measurements from upgoing S waves recoded at IRIS Champ temporary seismic network, denoted by black triangles. Red vectors represent splitting vectors. White arrows represent the Nazca APM.

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Figure 4-2. Rose histogram of fast directions from the entire local dataset. Dominant directions are approximately N40°E, N20°W, and N10°W-N20°E. Black arrow represents the mean direction of the dataset.

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Figure 4-3. Local fast polarization directions colored by direction. Red ~ Nazca APM; yellow ~ N40°E; blue ~ N20°W; green ~N10°W-20°E; black ~ other direction. Yellow dots represent cities in Chile.

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Figure 4-4. Schematic diagram of model from Long and Silver (2008), showing 3D mantle flow laterally around the slab dominating sub-slab flow, and the competing influence of 2D and 3D flow in the mantle wedge.

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Figure 4-5. Local S splitting measurements from MacDougall et al. (2012).

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CHAPTER 5 SUMMARY AND CONCLUSIONS

The dense aftershock distribution of the Maule 2010 Mw 8.8 earthquake, as recorded at broadband seismic stations of the International Maule Aftershock

Deployment (IMAD), provided an opportunity to study seismic properties of the Maule segment of the Nazca-South America subduction zone in detail. In this dissertation, I presented three studies aimed at investigating the rupture zone (33°S-38°S) of the 2010

Maule, Chile Mw 8.8 megathrust earthquake.

Using data from 197 local earthquakes (band-pass filtered from 0.7-20 Hz) recorded at the IRIS CHAMP IMAD stations, we calculated the differential seismic attenuation along 1,448 ray paths using two complementary spectral ratio P-S phase

-1 pair methods. The first method iteratively determines 400 individual Qs measurements and uncertainties for each P-S phase pair. The second method builds upon the first and stacks the spectra of each of the 400 measurements, yielding a single composite

-1 -1 spectrum from which we derived a single Qs value. From Qs we calculated t* values, using the IASPEI radial earth velocity model, which were incorporated into a 3D inversion to solve for attenuation structure in the Maule rupture zone above a depth of

140 km. We implemented a bounded linear inequality least squares inversion (0 < Qs <

-1 1200) to solve for Qs in individual model blocks with a constant starting model of 1/Qs

= 0, and solved for perturbations to the model using observed t* values and our velocity model. The inversion model clearly shows an east-dipping low attenuation

-1 features (Qs ~ 900), coincident with the location of the Nazca slab in this region. Above

-1 the Nazca slab, we image a region of higher attenuation (Qs ~400), which we interpret as the mantle wedge and overlying South American continental crust. Using the location

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of the Nazca slab from the SLAB1.0 model (Hayes et al., 2012), we calculated the mean

-1 Qs of the slab to be ~870 and that of the upper mantle wedge and crust to be ~505.

We interpret more localized features in the attenuation tomography model from

Chapter 2 to be related to regional geology. The large, low-attenuation feature in Figure

2-13 of our attenuation model is likely a decapitated Triassic marginal basin spreading ridge upper mantle structure imbricated into the overriding South American crust with no surface expression; incorporation of this structure into South American lithosphere is cogenetic with Triassic a-type granitoid magmatism (for example, the Cobquecura pluton), that outcrops sporadically in the Maule rupture region. Evidence for this interpretation includes regional geology north of ~36°S where Triassic intrusives are surficially exposed along the Coastal Cordillera. Also, the presence of fayalite in coeval granitic intrusives along the Coastal Cordillera suggests that these plutons were not emplaced in a subduction zone environment. We imaged additional low attenuation features at the surface (Figure 2-11, Figure 2-12), however, which could be related to the Triassic intrusives present at those locations, as is observable from comparison with the surface geologic units.

To determine the shear wave splitting within the Maule rupture zone from 57 teleseismic earthquakes recorded at IRIS CHAMP stations, we implemented the multichannel shear wave splitting intensity method of Chevrot (2000) using the SKS and

SKKS phases. This method takes advantage of that fact that, for typical Earth anisotropy, the slow component of the split shear wave is simply equal to the time derivative of the fast component, multiplied by a factor that is dependent on the splitting parameters,  and t. We found 386 suitable event-station pairs, yielding splitting

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measurements for 52 seismic stations. Split fast directions are largely trench-parallel

(~N30°E in the center of the Maule segment) with very few trench-normal measurements. Additionally, we do not observe many fast polarization directions that parallel the Nazca APM; instead, they trend mostly 40°-50° counterclockwise from

Nazca APM. We do not see a consistent spatial distribution of splitting delay times, with the majority of delay times ranging from 0.3 s – 2.8 s and two outliers at 3.45 s and 3.46 s. We interpret our measurements to be sampling mostly sub-slab mantle flow because the wavelengths of the SK(K)S phases used are too large to reliably resolve crustal anisotropy and fossilized fabrics in the subducting Nazca plate have likely been overprinted as a result of subduction. From these data, we do not see evidence for the commonly invoked 2D mantle flow model that elicits sub-slab entrained mantle flow to explain fast directions that parallel the APM of the subducting plate. Instead, our observations are consistent with the retrograde flow model of Russo and Silver (1994) which proposes lateral flow beneath the subducting Nazca oceanic plate to explain the trench-parallel flow along the Nazca-South American subduction zone, which has been widely observed by subsequent shear wave splitting studies (e.g., Polet et al. 2000;

Anderson et al., 2004).

We modified the Chevrot (2000) multichannel analysis method to incorporate local upgoing S phases to solve for shear wave splitting due to the anisotropy in the mantle wedge and overriding continental crust of the Maule rupture zone. Data from 132 aftershocks of the Maule event were used to calculate the splitting at 39 of the IRIS

CHAMP stations using 233 event-stations pairs. These measurements are much more variable than those from the teleseismic splitting dataset and have three dominant

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splitting directions of N20°W, N40°E, and N10°W-N20°E, with a subset of fast polarization directions oriented between N60°- 90°E. Fast polarization directions trending N60°-90°E are subparallel to the direction of Nazca APM, and we interpret these measurements to be representative of either (i) 2D corner flow in the mantle wedge or (ii) shearing directions in crustal rocks, consistent with the plate convergence direction. Splits oriented N20°W and N40°E are consistent with expected conjugate shearing directions, offset ~30° from orthogonal to the maximum compressive stress direction in the Maule segment, suggesting crustally sourced shear wave splitting. We expect anisotropy in crustal petrofabrics to largely influence the splitting measurements in this local dataset. Our third primary fast direction, however, reveals fast axes between

N10°W-N20°E, suggesting that we may be sampling mantle flow in the forearc, moving along-strike as a result of trench migration, as is the case for the sub-slab flow determined from our teleseismic study. These results from the Maule aftershocks are consistent with the model suggested by Long and Silver (2008) which they developed using data from 13 global subduction zones. Long and Silver (2008) proposed that mantle flow in the mantle wedge is trench-parallel near the subduction trench as a result of trench migration and then transitions to flow that is more consistent with the 2D corner flow model in which entrained mantle is coupled to the subducting lithosphere farther from the forearc.

In the future, we would like to incorporate both the teleseismic dataset and the local dataset of splitting intensity measurements together in a finite frequency inversion to constrain the anisotropy above, within, and below the subducting Nazca slab. The results we anticipate are highly relevant to understanding the geodynamics of

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subduction, as the fabric orientations in the upper 410 km of the Earth at subduction zones clearly result from plate interactions. Data like these would allow quantitative comparisons of detailed subduction zone deformation fabrics and structure along strike and possibly also between subduction zones.

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APPENDIX A SEISMIC ATTENUATION THEORY

A.1 Sources of Attenuation

Attenuation is the decrease in seismic wave amplitude as the wave propagates, and is a consequence of conservative elastic and non-conservative anelastic properties of the earth (Gordon and Davis 1968). Processes such as geometric spreading, scattering, and multipathing, can all result in attenuation of seismic energy (Stein and

Wysession, 2003). Geometric spreading begins immediately after the wave front starts propagating away from the source and is a function of distance from the source. As the wave front travels, the energy is distributed over a larger volume, causing the amplitude of the seismic wave to decrease. Another source of seismic attenuation is scattering, in which a seismic wave encounters a small-scale velocity or density perturbation and is forced to deviate from its current path. Scattered waves travel at varying speeds and can arrive at a recording seismic station at different times after the main phase. These multiple arrivals make it difficult to glean information from within the incoherent coda.

Yet another source for seismic attenuation is multipathing. Multipathing occurs when a seismic wave is assumed to have originated from a point source and then travels multiple paths to arrive at the seismic station. This breakup of the original wave front into multiple paths can be due to velocity or density heterogeneities along the wave path, similar to scattering.

A.2 Quantifying Attenuation

Seismic attenuation is an important parameter that can be used to quantify and physically characterize the Earth medium (Sato et al., 2000); it is quantified by a dimensionless quality factor, Q,

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1 E Q  (A-1) 2Emax where E is the amount of energy lost per cycle and Emax is the maximum elastic energy contained in a cycle (Knopoff 1964). As seen from Eq. A-1, the quality factor is inversely  proportional to attenuation per cycle, with high Q values corresponding to low attenuation and low Q values corresponding to high attenuation.

When we look at the decrease in seismic wave amplitude as a function of time in Eq. A-

2, we see that it is dependent on the quality factor.

0t 2Q (A-2) A(t)  A0e

Looking at the spectral ratio between the P and S phases for a particular frequency, 0, we can rewrite Eq. A-2 as Eq. A-3.

0 (ts t p ) A A0 2Q s  s e (A-3) A A p 0 p

When we take the natural logarithm of both sides of Eq. A-3 and rearrange, we see a linear relationship between frequency (0) and ln(As/Ap) and can solve for the quality factor, Q. To account for attenuation of body waves, we convolve the quality factor with an attenuation operator parameterized by t* (Lay and Wallace 1995). t t*  (A-4) Q

If we approximate the quality factor to vary as a function of depth, Q=Q(r), t* can be written as a path integral value as given by:

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N dt ti t*    (A-5)  Q Q path i1 i

As seen from Eq. A-4 and Eq. A-5, t* represents the total travel time divided by the path-averaged value of Q.

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APPENDIX B ANISOTROPY AND SHEAR WAVE SPLITTING THEORY

While most minerals are inherently anisotropic (Babuska and Cara 1991), random orientations of these crystals often nullify any anisotropy on the macroscopic scale observable by seismic waves. For anisotropy to be observable on the large scales of interest in seismology, deformation processes must either align the individual crystals of intrinsically anisotropic minerals in a lattice-preferred orientation (LPO) (e.g. Ribe

1989; Kaminski and Ribe, 2001; Karato, 1998a), or induce alignment of materials with elastically distinct properties, such as melt, in shape-preferred orientation (SPO)

(Kendall and Silver, 1996). The mantle is dominated by LPO of intrinsically anisotropic materials, although SPO may contribute in the shallowest mantle (e.g. Greve and

Savage, 2009) or in the D” region (e.g. Moore et al., 2004). The upper mantle is predominantly comprised of the mineral olivine, with a large single-crystal anisotropy of

~18% for shear waves (e.g. Mainprice, 2007), with an a-type fabric orientation (e.g.

McKenzie, 1979; Ribe and Yu 1991; Mainprice and Silver, 1992; Ismail and Mainprice,

1998) although occurrences of b-, c-, d- and e-type fabrics are also noted (e.g. Jung et al. 2006; Mainprice, 2007; Karato et al., 2008). Shear wave splitting, a phenomena in which a seismic wave incident upon an anisotropic region is split into two orthogonal waves, rotated from the fast polarization axis through an angle  and fast and slow quasi-shear waves offset by a time delay t, is a powerful tool in understanding the distribution of anisotropy in Earth material (Figure B-1) (Keith and Crampin, 1977;

Savage et al., 1990).

On a large-scale, because the LPO of the a-type olivine, which dominates upper mantle anisotropy under moderate stress and temperature conditions, orients fast axes

130

in upper mantle aggregates in the direction of maximum strain (e.g. Nicolas and

Christensen, 1987; Kaminski and Ribe, 2002), anisotropy constrained to the upper mantle from teleseismic shear wave splitting measurements is often attributed to either tectonic deformation processes (e.g. McKenzie, 1979) or the regional mantle flow regime (e.g. Tanimoto and Anderson, 1984; Vinnik et al., 1992; Anderson et al., 2004).

Olivine c-, d-, and e-type fabrics, while petrographically distinct from the a-type, yield fast polarization directions with the same orientation as produced by the a-type, i.e. paralleling the maximum strain direction (Jung et al., 2006; Kneller et al., 2008).

However, when olivine is subject to shear in hydrous, low temperature, high stress environments, b-type fabrics dominate, aligning the fast axes of olivine orthogonal to the maximum shear direction (e.g. Jung and Karato, 2001; Long and Becker, 2010), making it difficult to interpret shear wave splitting in terms of strain without understanding the thermal and mineralogical conditions of the subduction system. Many studies of shear wave splitting in subduction zones (e.g. Karato, 1995; Nakajima and Hasegawa, 2004;

Lassak et al, 2006; Kneller et al., 2008; Katayama et al., 2009; Bezacier et al., 2010) suggest that b-type olivine dominates the colder, more hydrous mantle nose in the forearc, whereas the a-, c-, d-, and e-type fabrics dominate the mantle wedge in the backarc.

Typically, receiver-side anisotropy is investigated using the teleseismic phases

PK(K)S and SK(K)S from events at distances between 87° <  < 150° from the recording station (Figure B-2). These phases are used because they propagate through the liquid outer core that converts shear wave energy to compressional energy, thereby removing any source-side shear wave splitting. When these waves leave the outer core

131

we assume that they are radially polarized, meaning that the direction of maximum particle excursion is in the radial plane. As a result, when we observe shear wave energy on the transverse component of the seismogram, it suggests leakage of energy from the radial component due to anisotropy along the ray path in the upper mantle. We can use shear wave splitting methods (e.g. Silver and Chan 1991, Wolfe and Silver

1998, Chevrot 2000) to determine the best  and t to correct for this energy leakage, minimizing the energy on the transverse component.

For local shear waves, we cannot assume the waves are radially polarized before splitting because they do not travel through the outer core. Instead, we need to determine the initial polarization of the event, which can be readily calculated from moment tensor solutions. To determine the shear wave splitting parameters, we can use the same methods as for teleseismic waves, but maximize the energy in the direction of the initial event polarization, not the radial direction.

It is important to recognize that if a wave is initially polarized in the same direction as either the fast or slow axes of anisotropy, then no splitting will be observed because only the fast or the slow shear wave will remain (Savage 1999). Such measurements that reveal no splitting are deemed nulls by the splitting methodologies of Silver and Chan (1991) and Wolfe and Silver (1998) and suggest that, if anisotropy is present, the fast direction is either parallel or perpendicular to the polarization of the shear waves, or the medium is isotropic (Savage 1999).

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Figure B-1. Incident shear wave pulse travels through anisotropic layer and is split yielding two transmitted pulses separated in time, one polarized in the fast direction that parallels the fast symmetry axis of the anisotropic material, and the other in the slow direction (Shearer 2009).

Figure B-2. Schematic of an SKS and SKKS ray path. The wave originates as an S wave and travels downwards until it reaches the core mantle boundary (CMB). At the CMB, the S wave is converted to a P wave due to the inability of S waves to travel through the liquid outer core. Leaving the core, the wave is converted back to an S wave. This new S wave is polarized only in the radial direction. As a result, shear wave splitting observed on an SKS phase is due to receiver side anisotropy. Modified from (Long and Silver 2009).

133

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BIOGRAPHICAL SKETCH

Megan Torpey was born in 1989 in New Jersey. She received a Bachelor of

Science degree in physics in May 2011 from The College of New Jersey, with a concentration in Earth science and a minor in French. In August 2011 she became a graduate student in the Department of Geological Sciences at the University of Florida.

She was awarded the Chateaubriand Fellowship by the Office of Science and

Technology at the French Embassy in 2015 and served as a research assistant in

France for the spring of 2015. She received her Ph.D. from the University of Florida in the spring of 2016.

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