The Salton Trough a Dissertation Submitted to the Department of Geophysics

Home , Garlock Fault, Peninsular ranges, Salton Buttes, Salton Trough, Sierra Nevada Batholith, Tanya Atwater

PASSIVE SEISMIC STUDY OF A MAGMA-DOMINATED RIFT: THE SALTON TROUGH

A DISSERTATION SUBMITTED TO THE DEPARTMENT OF GEOPHYSICS AND THE COMMITTEE ON GRADUATE STUDIES OF STANFORD UNIVERSITY IN PARTIAL FULFILLMENT OF THE REQUIREMENTS FOR THE DEGREE OF DOCTOR OF PHILOSOPHY

ii I certify that I have read this dissertation and that, in my opinion, it is fully adequate in scope and quality as a dissertation for the degree of Doctor of Philosophy.

(Simon Klemperer) Principal Adviser

I certify that I have read this dissertation and that, in my opinion, it is fully adequate in scope and quality as a dissertation for the degree of Doctor of Philosophy.

(Greg Beroza)

I certify that I have read this dissertation and that, in my opinion, it is fully adequate in scope and quality as a dissertation for the degree of Doctor of Philosophy.

(Norm Sleep)

Approved for the Stanford University Committee on Graduate Studies

iii Preface

This PhD thesis began with the aim to solely study the magmatic structure beneath Salton Trough transtensional rift basin in southernmost California using a temporary array of seismometers deployed across the basin and its margins. However, during the 2-year station deployment period, I added a large data set from permanent and other temporary seismometer arrays in aim to improve data coverage above the study region. As a result, I created a larger than planned Rayleigh-wave group-velocity model of the entire western U.S.A. and a shear-wave velocity model of southern California (See Chapter 2). In chapter 2, published in G-Cubed, I present new results and interpretations on the structure of the southern San Andreas Fault, the Peninsular Ranges batholith and southernmost Sierra Nevada batholith, in addition to this thesis’s main area of focus, the Salton Trough. The reader can find a detailed geological and geophysical background for the aforementioned regions in Chapter 2. The chapters in this thesis are divided by methodology. In chapter 2, I use ambient- noise tomography to create a 3D shear-velocity model of southern California’s crust and upper-mantle. The new 3D model reveals previously unimaged structure beneath the San Andreas Fault, Peninsular Ranges batholith, southern Sierra Nevada batholith and the Salton Trough, with implications on the geologic and tectonic history of the region. In chapter 3, published in Geology, I use shear-wave splitting of seismic waves, caused by anisotropic upper-mantle structure beneath the Salton Trough. I show evidence of a sharp rheological boundary in the upper-mantle beneath the Salton Trough region, at the southwest limit of the San Andreas Fault system, and infer the presence of upper-mantle melt-filled inclusions oriented parallel to transform motion on the San Andreas Fault. In the final chapter, I calculate receiver functions

iv on my array, and jointly invert them with my surface-wave dispersion curves (Chap- ter 2) to obtain a 2D shear-wave velocity model across the Salton Trough. I show improved imaging of the mantle upwelling beneath the rift.

v Acknowledgments

First and foremost I would like to thank my adviser Simon Klemperer, whose tireless enthusiasm and unfailing support made all of this work possible. Jesse Lawrence was a stalwart resource during my graduate studies, and I am grateful for his collaboration and advice over the years. The receiver-function stacking code used in Chapter 4 was principally of his writing. Greg Beroza and Norm Sleep, and George Thompson provided valuable comments during my annual reviews and Oral Defense. I thank Gary Fuis for sharing his SAF model and advice, Vicky Langenheim for sharing data and ideas, and Marty Grove for geological insights. I also thank Alex Kinsella (SURGE intern) and Monica Poletti (IRIS intern), my undergraduate interns, for improving my mentorship skills. And ﬁnally, the biggest thanks of all go to my parents, Yehuda and Merav, and my wife, Sivan, without whom none of this would have been possible.

vi Contents

Preface iv

Acknowledgments vi

1 Introduction 1

2 San Andreas Fault dip, Peninsular Ranges maﬁc lower crust and partial melt in the Salton Trough, Southern California, from ambient-noise tomography 6 2.1 Abstract ...... 7 2.2 Introduction ...... 9 2.3 Geological and Geophysical Background ...... 10 2.3.1 San Andreas Fault ...... 10 2.3.2 The Mesozoic Arcs ...... 10 2.3.3 The Salton Trough ...... 13 2.4 Data and Method ...... 15 2.4.1 Data ...... 15 2.4.2 Seismic Waveform Processing ...... 15 2.4.3 2-D Tomographic Inversion ...... 16

2.4.4 Inversion to VS(z) ...... 17 2.5 Rayleigh Group-Velocity Model ...... 19 2.6 Shear-Wave Velocity Model ...... 21 2.6.1 San Andreas Fault ...... 25 2.6.2 The Mesozoic Batholiths: Peninsular Ranges and Sierra Nevada 31

vii 2.6.3 Salton Trough ...... 39 2.7 Summary ...... 47 2.8 Appendix: Calculation of Pore Geometry and Partial Melt ...... 48 2.9 Acknowledgments ...... 49

3 Rapid variation in upper-mantle rheology across the San Andreas fault system and Salton Trough, southernmost California 50 3.1 Abstract ...... 51 3.2 Introduction ...... 51 3.3 Analysis of Shear-Wave Splitting Measurements ...... 53 3.4 Results ...... 55 3.5 Discussion ...... 56 3.6 Conclusions ...... 59 3.7 Acknowledgments ...... 59

4 Imaging upper-mantle upwelling beneath the Salton Trough using joint inversion of receiver functions and surface wave dispersion curves 60 4.1 Abstract ...... 62 4.2 Introduction ...... 62 4.3 Data and Method ...... 62 4.4 Results ...... 62 4.5 Discussion ...... 62 4.6 Conclusions ...... 62 4.7 Acknowledgments ...... 62

A Data Acquisition 63

viii List of Tables

ix List of Figures

1.1 A: Paciﬁc (PAC)-North America (NAM) plate boundary (black: transform, orange: spreading center), showing location of B (black rectangle). B: Shaded-relief topography overlain with the broadband stations of the Salton Seismic Imaging Project (stars). CM-Chocolate Moun- tains; PRB-Peninsular Ranges Batholith; ST-Salton Trough. Red lines show major faults of the San Andreas fault system (SAFs) (AF- Algodones, EF-Elsinore, IF-Imperial, LSF-Laguna Salada, SAF-San Andreas, SHF-Sand Hills, SJF-San Jacinto) (Plesch et al., 2007). . . .2

x 2.1 (a) Map of stations used for ambient noise cross correlations, colored by recording epochs. Physiographic provinces from USGS (1946). Rect- angle shows area of (b) & (c). GOC: Gulf of California. Figure 2.1b shows Geographical and geological features. Broad black-dashed line: Compositional and geochemical boundary between the western Penin- sular Ranges Batholoith (PRB) and the eastern PRB from Todd et al. (1988). Red lines: fault traces after Plesch et al. (2007). AF: Algodones Fault; BSZ: Brawley Seismic Zone; CM: Chocolate Mountains; CP: Cerro Prieto volcano; CR: Coast Ranges; EF: Elsinore Fault; GF: Gar- lock Fault; GV: Great Valley; IF: Imperial Fault; IV: Imperial Valley; KCF: Kerns Canyon Fault; LSB: Laguna Salada Basin; LSF: Laguna Salada Fault; OVF: Owens Valley Fault; ePRB and wPRB: east and west Peninsular Ranges Batholith; SAF: San Andreas Fault; SB: Salton Buttes; SBM: San Bernardino Mountains; SGM: San Gabriel Moun- tains; SGP: San Gorgonio Pass (black dot); SHF: Sand Hills Fault; SJF: San Jacinto Fault; SNB: Sierra Nevada Batholith WTR: West- ern Transverse Ranges. Controlled source proﬁles: NB82: (Nava and Brune, 1982); PM96: (Parsons and McCarthy, 1996); H13:(Han et al., 2013). Figure 2.1c shows locations of proﬁles in Figures 2.6– 2.11: San Andreas Fault: AA0–II0 (red); Peninsular Ranges: (PR)aa0–cc0 (green); Sierra Nevada: (SN)aa0–bb0 (blue); Salton Trough: (ST)aa0–cc0 (yellow). Dipping SAF model (Fuis et al., 2012) colored by depth. . . . .8

2.2 Schematic of 13 starting models used for each 1-D proﬁle of VP , VS and ρ extracted from CVM-H11.9, created by increasing and decreasing the velocity up to 15%...... 18 2.3 (left) Comparison of our Rayleigh group-velocity model to (right) Moschetti et al. (2007) showing deviation from the mean group velocity (U) at 14 s. Dashed-line rectangle: area of the model we inverted for shear-wave velocity. Solid thin lines: state and physiographic boundaries after USGS (1946)...... 20

xi 2.4 Maps of our Rayleigh group-velocity model in southern California, at periods of 8, 16, 24, and 40 s, shown as deviation from the mean group velocity (U) of the entire western U.S. model for that period, with color range varying between maps due to the large variation in velocity perturbations. Annotations as in Figure 2.1b...... 22

2.5 Depth slices through our ﬁnal Vs model at 7, 20 and 42 km depth. . . 24 2.6 Caption on previous page...... 27 2.6 ...... 28 2.6 ...... 29 2.7 (a) Depth to the shallowest 3.6 km/s isovelocity surface, showing deepening (lower velocities) NE of the SAF and NE of the boundary between wPRB and ePRB. (b) Depth to the shallowest 3.8 km/s isovelocity surface, showing deepening (lower velocity) NE of the PRB and E of the prebatholithic suture in the SNB (see Figure 2.10). Fault traces and physiographic divisions as in Figure 2.1. Depth contours at 5km interval. 32

2.8 West-east proﬁles across the PRB through our VS ﬁnal model, as shown in Figure 2.1c. Figure annotation as in Figure 2.6. Dashed line: compositional boundary between wPRB and ePRB, extrapolated to depth at a 45◦ angle (cf. Langenheim et al., 2014)...... 33 2.9 Comparison along PRb–PRb0 (Figure 2.8b) between our model (a), CVM-H11.9 (initial model; b), and CVM-S4.26 (c). (d) Standard deviation of our model. Figure annotation as in Figure 2.8. See also supporting information Figures S51–S53...... 34 2.10 Caption on previous page...... 38

xii 2.11 VS profiles through our final model in the Salton Trough region. (a) STa–STa0 along the axis of the trough. (b) STb–STb0 perpendicular to the axis of the Salton Trough. (c) STc–STc0 north-south profile that show both low-velocity zones in the upper mantle described in the text. Vertical dashed lines: intersection of profiles. Note very different vertical exaggeration of part (Figure 2.11a). Standard deviation of our model, and comparison with CVM-H11.9 and CVM-S4.26 shown in supporting information Figures S56–S58. Abbreviations from Figure 2.1b...... 40 2.12 Horizontal slice at 42 km depth showing low-velocity zones in the upper mantle beneath the Salton Trough. Green lines: coast and state borders. Purple lines: surface trace of faults as in Figure 2.1b. Dot- patterned boxes: active spreading centers after Elders et al. (1972). Other annotations as in Figure 2.1b...... 41

xiii 2.13 Plot of laboratory and ﬁeld measurements of VP versus VS to estimate composition and volume of partial melt beneath the Salton Trough (a) in the crust and (b) in the upper mantle. Field measurements: gray- shaded rectangles show the observed range of velocities constrained by

our model (VS) and active-source data (VP ) Fuis et al. (1984); Han et al. (2013); Parsons and McCarthy (1996). Red lines and arrows show the

upper limit or range of VP for lower crust and upper mantle from these three separate seismic experiments. In the crust, LVZ: lower-crustal low-velocity zone; HVZ: midcrustal high-velocity zone. In the mantle, Lid: normal velocity upper-mantle lid; LVZ: upper-mantle low velocity zone. Laboratory measurements: Squares Christensen (1996) and diamonds Wang et al. (2013a) are colored yellow for lower temperature measurements (20 or 25◦C at 600 MPa in Figure 2.13a, 600◦C at 1000 MPa in Figure 2.13b) and orange for higher temperature (600◦C at 600 MPa in (a) and 1000◦C at 1000 MPa in (b)). Rock types: AGR: Anorthositic granulite; AMP: Amphibolite; ANO: Anorthosite; DIA: Diabase; DIO: Diorite; DUN: Dunite; FGR: Felsic granulite; GAB: Gabbro; MGR: Maﬁc granulite; PER: Peridotite; PYX: Pyroxenite. Error bars show the mean error for the data from Christensen et al., 1996Christensen (1996) and Wang, Q. et al., 2013Wang et al. (2013a). In both Figures 2.13a and 2.13b, we show the eﬀect of melt on seismic wavespeed for a single rock type (gabbro in the lower crust, peridotite in the upper mantle). Lines of colored squares originating at these rock types (GAB and PER) show volume of partial melt along curves of con-

stant dlnVS/dlnVP (tilted numbers, 1.0, 1.6, 2.25) calculated for these rock types. A line of colored squares that passes through both dark gray (HVZ, Lid) and light gray (LVZ) shaded-rectangles corresponds to a single rock composition that could represent both the high- and low-velocity lower crust, or both the lid and low-velocity upper mantle, with diﬀerent melt fractions controlling the change in velocity. Thin-

dotted lines are lines of constant VP /VS...... 45

xiv 3.1 A: Pacific (PAC)-North America (NAM) plate boundary (black: transform, orange: spreading center), showing location of B and C (black rectangle), and shear-wave splitting (SWS) fast polarizations further south in Mexico (Long, 2010) using length scale and color scale as in Figure 3.1. B: Shaded-relief topography overlain with our seismic stations (stars), other permanent stations shown in Figure 3.2 (triangles), and line of section in Figure 3.2 (white line). BRP-southern Basin and Range province; CMR-Chocolate Mountains Region; PRB-Peninsular Ranges Batholith; ST-Salton Trough. Dot-patterned orange boxes show active spreading centers (Elders et al., 1972). Strike-slip focal mechanism: 4 April 2010 El Mayor-Cucapah earthquake. Thick red lines show major faults of the San Andreas fault system (SAFs) (AF- Algodones, CPF-Cerro Prieto, EF-Elsinore, IF-Imperial, LSF-Laguna Salada, SAF-San Andreas, SHF-Sand Hills, SJF-San Jacinto). Thin red lines show southeast extension of Eastern California Shear Zone (ECSZ) (Darin and Dorsey, 2013). C: SWS fast polarizations with lengths proportional to delay time (colored bars). This study: white circles (open circles if no “good” or “fair” results were available); other SWS results: Wüstefeldet al. (2009); Liu et al. (2014). (2014; only ‘A’ quality data, weighted by their standard deviation). Color of the bars indicates difference in degrees from trend of SAFs. Black arrows are plate-motion vectors relative to the fixed-hotspot frame (Gripp and Gordon, 2002). Red line is 55 km depth contour of lithosphere- asthenosphere boundary (Lekic et al., 2011). Thin black lines are fault traces (Plesch et al., 2007). Copyright (2016) Geological Society of America...... 52

xv 3.2 A: Our shear-wave splitting results and 10 previous measurements (Fig. 3.1B), projected onto the white line in Figure 3.1B. Thick (and thin) solid colored lines are least-squares linear ﬁts to split directions (and split times) for each group of stations (Peninsular Ranges Batholith [PRB]; Salton Trough [ST]; Chocolate Mountains Region [CMR]; southern Basin and Range province [BRP]). Error bars are quality-weighted standard deviation of polarization direction; circle diameters are proportional to split time. Thin horizontal magenta lines show Paciﬁc and North America absolute (PAC(HS), NAM(HS)) and relative (PAC(NA)) plate motions. Thick gray bars show Far- allon (FAR) motion with respect to the North America plate, and strike of San Andreas fault system (EF-Elsinore, SJF-San Jacinto, IF- Imperial, SHF-Sand Hills, SAF-San Andreas fault; ECSZ-Eastern Cal- ifornia Shear Zone). B: S-velocity model (color-scale) (Barak et al., 2015) overlain with lithosphere-asthenosphere boundary (LAB) from Sp receiver functions (white circles: Lekic et al. (2011)). Black lines are example SKS ray-paths, using the mean event azimuth, distance and depth. C: Tectonic cartoon. Numbers above fault symbols are maximum estimates of geodetic fault slip rates, mm/yr (Chuang and Johnson, 2011; Field et al., 2013). Yellow ovals are aligned melt channels. Curved arrows show upwelling mantle. Dashed lines are inferred shear zones. ⊗ and are block motions into and out of the page, size crudely scaled to velocity with respect to the EF. Copyright (2016) Geological Society of America...... 54

xvi Chapter 1

Introduction

The Salton Trough (ST) is a large transtensional rift basin, which lies in the transition zone between the Gulf of California oblique-rift system and the right-lateral San Andres Fault (Figure 1.1). Although extensively studied in the past, still very little is known about the total volume of intrusion into the crust, or the magma distribution within and beyond the rift margins. Previous efforts to seismically image the crustal structure in southern California include body-wave and surface-wave tomography (e.g., Yang and Forsyth, 2006; Tape et al., 2010; Schmandt and Humphreys, 2010; Allam and Ben-Zion, 2012), receiver-function analysis (e.g., Yan and Clayton, 2007a; Lekic et al., 2011) and controlled-source studies (e.g., Fuis et al., 1984, 2003; Han et al., 2016; Persaud et al., 2016). However, these event-based methods are limited by the temporal and spatial distribution of their sources and receivers, rendering the 3-D structure of the lower crust and upper mantle poorly resolved. In recent years, ambient-noise surface-wave tomography (ANT) has emerged as a useful tool for crustal imaging (e.g., Shapiro et al., 2005; Zigone et al., 2015). In contrast to traditional methods, ANT does not depend critically on the spatial and temporal distribution of earthquakes and can provide homogeneously distributed information about shear-wave velocities through the entire crust and uppermost mantle (Shapiro et al., 2005). In southern California especially, ANT benefits from the abundance of broadband network stations and noise sources (Pacific ocean waves and local seismicity). Southern California’s subsurface structure has previously been imaged with ANT in both local (e.g., Shapiro et al., 2005; Sabra et al., 2005; Zigone et al., 2015)

1 CHAPTER 1. INTRODUCTION 2

−117˚ −116˚ −115˚ 34˚ 34˚ NAM

S A F PAC

S JF CM

E F SH F AZ 33˚ CA 33˚ PRB

IF USA A F Mexico ST

L B SF A

−117˚ −116˚ −115˚

Figure 1.1: A: Pacific (PAC)-North America (NAM) plate boundary (black: transform, orange: spreading center), showing location of B (black rectangle). B: Shaded-relief topography overlain with the broadband stations of the Salton Seismic Imaging Project (stars). CM-Chocolate Mountains; PRB-Peninsular Ranges Batholith; ST-Salton Trough. Red lines show major faults of the San Andreas fault system (SAFs) (AF-Algodones, EF-Elsinore, IF-Imperial, LSF-Laguna Salada, SAF- San Andreas, SHF-Sand Hills, SJF-San Jacinto) (Plesch et al., 2007). CHAPTER 1. INTRODUCTION 3 and regional (e.g., Moschetti et al., 2007) studies. Full-waveform inversion of both body-waves from regional earthquakes and ambient-noise derived correlograms has also been done (Lee et al., 2014). Seismic imaging can further improve by jointly inverting receiver functions (e.g., Langston, 1977) and high frequency surface wave dispersion data (e.g., Juliàet al., 2000). Receiver functions can image large-scale discontinuities, such as the Moho, that are not well-resolved by the ambient noise tomography alone (e.g., Shen, 2013). Despite Southern Californias (USA) complex tectonic history and active plate boundary oriented northwest-southeast (e.g., Dickinson, 2008, 2009; Barak et al., 2015) (Chapter 3, Fig. 1A), previous studies of shear-wave splitting in Southern California (e.g., Polet and Kanamori, 2002; Kosarian et al., 2011) show a nearly uniform fast axis of anisotropy oriented approximately west-east (Chapter 3, Fig. 1C) interpreted as due to inherited North America plate motion (Kosarian et al., 2011), or to pre-late Cenozoic compression (Polet and Kanamori, 2002), or to mantle flow around the southern edge of the subducting Gorda slab (Zandt and Humphreys, 2008). The same general west-east pattern continues to the southern tip of Baja California (Mexico), west and east of the Gulf of California rift margins (Long (2010); Chapter 3, Fig. 1A). Recent three-dimensional (3-D) tomographic inversion of shear- wave splitting measurements, despite the previous scarcity of data in southernmost California, has begun to suggest complicated structure, including a northwest fast axis of anisotropy in the Salton Trough (ST) (Monteiller and Chevrot, 2011) where Gulf of California ocean spreading propagates into continental crust along the San Andreas fault (Elders et al., 1972). In January 2011, I deployed 42 broadband seismometers across the Salton Trough, provided by the PASSCAL pool and Caltech, as part of the Salton Seismic Imaging Project (SSIP) (Figure 1.1B). 36 of the 42 stations were deployed at 5-10 km intervals, along a line perpendicular to the rift axis. The other six stations, including two in Mexico, were deployed off-line to provide areal coverage. This array scheme was chosen to allow for multiple crossing ray-paths in the crust and upper mantle, important both for Common-conversion-point migration of receiver functions and ambient noise tomography. Stations were nominally deployed for 18 months in order to provide a good ensemble of large earthquakes with good azimuthal distribution, necessary for CHAPTER 1. INTRODUCTION 4 receiver function analysis. Actual recording time varied between 6 to 24 months. Our main line coincides with older and recent refraction profiles Fuis et al. (1984); Parsons and McCarthy (1996) and the more recent SSIP controlled-source profile, acquired in March 2011, both across and along the STRose et al. (2013); Persaud et al. (2016). The controlled-source part of the SSIP was lead by John Hole, Joanne Stock and Gary Fuis. In this work, I use ambient-noise tomography to create a new Rayleigh-wave group- velocity model of most of western U.S.A. (Chapter 2). Consequentially, I create a shear-wave velocity model within the southern California region by inverting the group-velocity dispersion curves for the shear-wave velocity structure. I use 4 years of data recorded on 849 broadband stations, vastly more than previous studies, including my station-array across the Salton Trough and other campaign stations in both Mexico and the United States. To study the anisotropic structure beneath the Salton Trough, I analyze shear-wave splitting measurements on my array of seismometers (Chapter 3). Analysis of teleseismic shear-wave splitting is a standard tool to study upper-mantle anisotropy created by strain-induced lattice-preferred orientation of minerals or by preferentially oriented melt-filled inclusions, hence also changes in rheology (e.g., Savage, 1999, and references therein). Finally, I use receiver functions calculated at my stations in a joint inversion with my previously calculated group- velocity dispersion curves, to improve imaging of the Moho and magmatic structure beneath the Salton Trough. My new shear-wave velocity model of southern California shows a mean lower-crust and upper-mantle wavespeeds (3.6 km/s at 20km, 4.2 km/s at 40km) that are low by global standards. Across the San Andreas Fault, southeast of San Gorgonio Pass, we observe vertical to steeply dipping lateral velocity contrasts that extend beneath the Moho. Beneath the western Peninsular Ranges batholith and westernmost southern Sierra Nevada batholith, our model shows relatively high shear-velocities (≥3.8 km/s) in the lower crust that I interpret as the mafic roots of the overlying arc. Relatively high-velocity upper-mantle (up to ∼4.5 km/s) may be part of the intact arc, or possibly a remnant of the Farallon plate. Beneath the Salton Trough, I image zones of low shear-velocity in the lower crust and upper mantle which permit up to ∼4.5% melt in the lower crust and up to ∼6% melt in the upper mantle, depending on the CHAPTER 1. INTRODUCTION 5 assumed composition and pore geometry. The results preclude the existence of older continental crust beneath the Salton Trough and support the creation of new crust beneath the Salton Trough. The shearwave splitting analysis shows systematic lateral variations in anisotropy between different geologic regions across the Pacific-North America plate boundary at the Salton Trough. The data suggest new constraints on the underlying upper-mantle rheology, and implications for active tectonics and associated earthquake hazard. The new 2D shear-wave velocity, obtained from the joint inversion of receiver functions and ambient noise dispersion curves, has a much larger depth extent (150 km), and reveals a more detailed structure of the magmatic plumbing in the upper mantle beneath the Salton Trough region. Chapter 2

San Andreas Fault dip, Peninsular Ranges maﬁc lower crust and partial melt in the Salton Trough, Southern California, from ambient-noise tomography

An edited version of this paper was published in G-Cubed by AGU. Copyright (2015) American Geophysical Union: Barak, S., Klemperer, S. L. & Lawrence, J. F. San Andreas Fault dip, peninsular ranges maﬁc lower-crust and partial melt in the Salton Trough, Southern Califor- nia, from ambient-noise tomography. Geochemistry, Geophys. Geosystems n/an/a (2015). doi:10.1002/2015GC005970. To view the published open abstract, go to http://dx.doi.org and enter the DOI.

Supplementary Material for this paper is available at the Stanford Digital Repos- itory: https://purl.stanford.edu/df776kg9323

6 CHAPTER 2. BARAK ET AL., G-CUBED 2015 7

2.1 Abstract

We use ambient-noise tomography to improve CVM-H11.9, a community velocity model of southern California. Our new 3D shear-velocity model with 0.05◦x0.05◦ lateral and 1 km vertical blocks reveals new structure beneath the San Andreas Fault (SAF), Peninsular Ranges batholith (PRB), southern Sierra Nevada batholith (SNB) and the Salton Trough (ST). We use 4 years of data recorded on 849 broadband stations, vastly more than previous studies and including our own broadband Salton Seismic Imaging Project, a 40 station transect across the ST, as well as other campaign stations in both Mexico and the United States. Mean lower-crust and upper-mantle wavespeeds (3.6 km/s at 20km, 4.2 km/s at 40km) are low by global standards. Across the SAF, southeast of San Gorgonio Pass, we observe vertical to steeply dipping lateral velocity contrasts that extend beneath the Moho. Beneath the western PRB and westernmost southern SNB, we observe relatively high shear-velocities (≥3.8 km/s) in the lower crust that we interpret as the maﬁc roots of the overlying arc. Relatively high-velocity upper-mantle (up to ∼4.5 km/s) may be part of the intact arc, or possibly a remnant of the Farallon plate. Beneath the ST, we observe zones of low shear-velocity in the lower crust and upper mantle which permit up to ∼4.5% melt in the lower crust and up to ∼6% melt in the upper mantle, depending on the assumed composition and pore geometry. Our results preclude the existence of older continental crust beneath the ST and support the creation of new crust beneath the ST. CHAPTER 2. BARAK ET AL., G-CUBED 2015 8

Figure 2.1: (a) Map of stations used for ambient noise cross correlations, colored by recording epochs. Physiographic provinces from USGS (1946). Rectan- gle shows area of (b) & (c). GOC: Gulf of California. Figure 2.1b shows Geo- graphical and geological features. Broad black-dashed line: Compositional and geochemical boundary between the western Peninsular Ranges Batholoith (PRB) and the eastern PRB from Todd et al. (1988). Red lines: fault traces after Plesch et al. (2007). AF: Algodones Fault; BSZ: Brawley Seismic Zone; CM: Choco- late Mountains; CP: Cerro Prieto volcano; CR: Coast Ranges; EF: Elsinore Fault; GF: Garlock Fault; GV: Great Valley; IF: Imperial Fault; IV: Impe- rial Valley; KCF: Kerns Canyon Fault; LSB: Laguna Salada Basin; LSF: La- guna Salada Fault; OVF: Owens Val- ley Fault; ePRB and wPRB: east and west Peninsular Ranges Batholith; SAF: San Andreas Fault; SB: Salton Buttes; SBM: San Bernardino Mountains; SGM: San Gabriel Mountains; SGP: San Gor- gonio Pass (black dot); SHF: Sand Hills Fault; SJF: San Jacinto Fault; SNB: Sierra Nevada Batholith WTR: Western Transverse Ranges. Controlled source proﬁles: NB82: (Nava and Brune, 1982); PM96: (Parsons and McCarthy, 1996); H13:(Han et al., 2013). Fig- ure 2.1c shows locations of proﬁles in Fig- ures 2.6– 2.11: San Andreas Fault: AA0– II0 (red); Peninsular Ranges: (PR)aa0– cc0 (green); Sierra Nevada: (SN)aa0–bb0 (blue); Salton Trough: (ST)aa0–cc0 (yellow). Dipping SAF model (Fuis et al., 2012) colored by depth. CHAPTER 2. BARAK ET AL., G-CUBED 2015 9

2.2 Introduction

Previous efforts to seismically image the crustal structure in southern California include body-wave and surface-wave tomography (e.g., Yang and Forsyth, 2006; Tape et al., 2010; Schmandt and Humphreys, 2010; Allam and Ben-Zion, 2012), receiver- function analysis (e.g., Yan and Clayton, 2007a; Lekic et al., 2011) and controlled- source studies (e.g., Fuis et al., 1984, 2003; Rose et al., 2013). However, these event- based methods are limited by the temporal and spatial distribution of their sources and receivers, rendering the 3-D structure of the lower crust and upper mantle poorly resolved. In recent years, ambient-noise surface-wave tomography (ANT) has emerged as a useful tool for crustal imaging (e.g., Shapiro et al., 2005; Zigone et al., 2015). In contrast to traditional methods, ANT does not depend critically on the spatial and temporal distribution of earthquakes and can provide homogeneously distributed information about shear-wave velocities through the entire crust and uppermost mantle (Shapiro et al., 2005). In southern California especially, ANT benefits from the abundance of broadband network stations and noise sources (Pacific ocean waves and local seismicity). Southern California’s subsurface structure has previously been imaged with ANT in both local (e.g., Shapiro et al., 2005; Sabra et al., 2005; Zigone et al., 2015) and regional (e.g., Moschetti et al., 2007) studies. Full-waveform inversion of both body-waves from regional earthquakes and ambient-noise derived correlograms has also been done (Lee et al., 2014). In this study, however, unlike previous ones, we use all available permanent network stations and USArray stations as well as campaign arrays (Figure 2.1a). These campaign arrays include our own Broadband Salton Seismic Imaging Project (bbSSIP) deployed across the Salton Trough, part of the larger SSIP program (Hole et al., 2011; Rose et al., 2013). In addition, while many previous ANT studies focused on methodology, this paper takes the opposite approach by adding more detailed geologic interpretations. We use the ambient-noise records to obtain a Rayleigh-wave group-velocity model of southern California between periods of 4 and 40 s. We then estimate the 3-D shear velocity (VS) distribution of the crust and uppermost mantle using CVM-H11.9 (Plesch et al., 2011) as a starting point. Finally, we describe and interpret prominent velocity anomalies observed in our group-velocity model and final shear-wave model. CHAPTER 2. BARAK ET AL., G-CUBED 2015 10

2.3 Geological and Geophysical Background

2.3.1 San Andreas Fault

The San Andreas Fault (SAF), a major transform plate boundary, is one of the most studied faults in the world, yet there is still much debate about its geometry and depth extent. Seismic images across the SAF in southern California are few. Although the SAF is commonly assumed to be vertical to steeply dipping (e.g., Plesch et al., 2007), a recent model for the geometry of the SAF in southern California describes the fault as having a propeller shape, dipping NE southeast of the Mojave Desert, turning vertical in the Mojave Desert, and then dipping SW from the Big Bend northward (Figure 2.1c) (Fuis et al., 2012). This dipping SAF model is based on modeling of potential-field data, active-source imaging, and seismicity, where available, interpolated along the SAF throughout southern California. More data and higher resolution are needed to confirm or refute this new propeller model. The depth extent of the SAF as a narrow fault or shear zone is also uncertain. Although the seismogenic thickness of the SAF is largely limited to depths of 11–20 km in southern California (Smith-Konter et al., 2011), evidence of a Moho offset (Zhu, 2002; Yan and Clayton, 2007b) or a downward bulge (e.g., Fuis et al., 2012) in the Mojave section of the fault suggests that the fault extends through the crust. Fuis et al. (2012) also project the SAF to the northern edge of a high-velocity body in the mantle that may represent the plate boundary to depths of ∼200 km.

2.3.2 The Mesozoic Arcs

The Peninsular Ranges

The Peninsular Ranges Batholith (PRB) is part of the Mesozoic magmatic arc formed above the subducting Farallon plate along the western margin of North America (Dickinson, 2009; Fenby and Gastil, 1991; Jennings et al., 1977). The PRB extends nearly 1200 km from south of the Transverse Ranges in the United States almost to the tip of the Baja California peninsula in Mexico (Langenheim et al., 2014). The PRB is divided into western (wPRB) and eastern (ePRB) zones with distinct ages, chemistry, and geophysical properties. Exposure level increases considerably from west to east CHAPTER 2. BARAK ET AL., G-CUBED 2015 11 across the PRB, with most exposures having crystallization pressures of 1–3 kb (∼4– 11 km) in the wPRB, compared to 3–6 kb (∼11–23 km) beneath the ePRB (Ague and Brimhall, 1988). Rocks range from olivine gabbro through felsic granite, but tonalite is the most abundant rock type across the entire PRB (Silver and Chappell, 1988). The wPRB consists of older, Jurassic-Early Cretaceous arc rocks, dominantly 125–100 Ma plutons with relatively primitive island-arc geochemical affinities. In contrast, the ePRB is dominated by a large-volume, 95±3 Ma tonalite-trondjemite- granodiorite suite formed from deep melting with garnet in its source (Gastil et al., 1975; Kimbrough et al., 2014; Silver and Chappell, 1988; Todd et al., 1988). The wPRB has abundant gabbro, plutons with magnetite±ilmenite oxides, no S-type granites, and δ18O '6.0–8.5. In contrast, the ePRB has very little gabbro, ilmenite as the dominant or only opaque oxide, abundant S-type granites and δ18O '9–11 (Silver and Chappell, 1988; Todd et al., 1988). The compositional boundary coincides locally with shear zones with steeply east-dipping metamorphic foliation (Johnson et al., 1999; Thomson and Girty, 1994). Gravity, magnetic, and seismic models suggest that the compositional boundary dips east to mid-to-lower crustal depths, separating relatively mafic, dense magnetic rocks of the wPRB from more silicic, less dense and weakly magnetic rocks of the ePRB (Magistrale and Sanders, 1995; Langenheim et al., 2014). Most authors have suggested that the compositional boundary is the result of juxtaposing oceanic crust to the west of transitional or continental crust (e.g., Sed- lock, 2003). “Single-arc” models propose that the batholith was emplaced across a preexisting transition between oceanic and continental lithosphere at the continental margin (e.g., Thomson and Girty, 1994; Todd et al., 2003). “Composite-arc” models that interpret the western zone as an accreted island arc that collided with North America (e.g., Johnson et al., 1999) now seem unlikely based on a detrital zircon population of North American affinity in the marine strata outboard of the wPRB (Kimbrough et al., 2014). Though the wPRB and ePRB originated within the same marginal arc system, the wPRB is a classic arc formed above older oceanic or arc basement, whereas the ePRB formed as a magmatic flareup during an abrupt east- ward relocation of the batholith coinciding with major underplating, devolatilization, and melting of the underthrust accretionary complex (Catalina schist) (Grove et al., CHAPTER 2. BARAK ET AL., G-CUBED 2015 12

2008). Knowledge of the 3D crustal velocity structure of the PRB has been limited to the upper 20 km (Magistrale and Sanders, 1995; Hauksson, 2000; Zigone et al., 2015), though an average lower-crustal VS of 3.8–4 km/s was estimated from a 2D seismic refraction proﬁle along the wPRB (Nava and Brune, 1982) (Figure 2.1b). The crustal thickness of the PRB varies from ∼33 km at the Paciﬁc coast increasing gradually toward the east within the wPRB to a maximum of ∼40 km (Lewis et al., 2001). From this maximum depth the crust thins rapidly toward the east across the topographically high ePRB to a crustal thickness of ∼15–18 km at the margin of the Gulf of California. Although it is possible that the compositional boundary between the wPRB and ePRB extends below 20 km (Langenheim et al., 2014), the boundary has not been previously imaged seismically in the lower crust.

Southern Sierra Nevada

Prior to the initiation of the SAF, the Sierra Nevada Batholith (SNB) and the PRB were part of the same arc. As summarized by Saleeby (2003), the Cretaceous SNB is very similar to the PRB, generated and emplaced into a metamorphic framework across a pre-batholithic suture (PBS) between Paleozoic oceanic lithosphere to the west and North American Proterozoic lithosphere to the east. The lithologic expres- sion of this transverse compositional gradient, as in the PRB, is a western domain rich in mafic and tonalitic rocks, and an eastern domain predominately granodioritic in composition. The SNB also displays a pronounced transverse age gradient with mafic magmatism commencing at ∼130 Ma in the west and felsic magmatism ending in the east by ∼80 Ma. In the southern SNB, as in the PRB, exposure depth increases from west-to-east; but also toward the Garlock Fault along which the SNB is upturned (Ague and Brimhall, 1988). By analogy with the PRB, we use western SNB (wSNB) to relate to the more mafic, older region west of the PBS, and eastern SNB (eSNB) for the more felsic and younger batholith. The coupled Sierra Nevada and Great Valley basin are recognized as a relatively rigid block (Sierra Nevada microplate) moving within the San Andreas-Walker Lane dextral system (Saleeby et al., 2013). Fliedner et al. (2000) showed that crustal thickness is, as in the PRB, somewhat greater (≤42 km) beneath the western part CHAPTER 2. BARAK ET AL., G-CUBED 2015 13 of the SNB than beneath the highest topography to the east. Gravity modeling requires increased densities in the crust or a lithospheric root or both, offset to the west of the exposed SNB (Fliedner et al., 1996), that is in the same location as a high-velocity upper-mantle anomaly (the “Isabella anomaly”) observed in many studies (e.g., Jones et al., 2014). The Isabella anomaly is commonly interpreted as the result of delaminated continental lithosphere or a convective drip beneath the Sierra Nevada (e.g., Zandt et al., 2004), although others suggest that the anomaly represents a fragment of the Farallon slab (Biasi and Humphreys, 1992; Wang et al., 2013b).

2.3.3 The Salton Trough

The Salton Trough (ST) in southern California is the result of oblique extension linking the Gulf of California (GOC) rift system, and the San Andreas Fault (SAF) transform system (e.g., Elders et al., 1972; Stock and Hodges, 1989). At ∼12 Ma (e.g., Stock and Hodges, 1989) or even as early as ∼17–20 Ma (e.g., Nicholson et al., 1994), when subduction ceased off the southwest margin of Baja California, rifting began in the Salton Trough segment. From ∼12 to ∼6 Ma, oblique transtension was partitioned between ENE extension (similar to Basin-and-Range extension) in the GOC and dextral slip on the Tosco-Abreojos fault offshore and west of Baja California (e.g., Stock and Hodges, 1989). Since at least ∼6 Ma, transform motion stepped inland, transferring Baja California and southwestern California to the Pacific plate. Today, large dextral strike-slip faults that carry most of the Pacific-North American plate motion cut through the basin. The SAF bounds the ST on the northeast and, south of the Salton Sea, plate motion is thought to be transferred into a series of small spreading segments connected by the Imperial and Cerro Prieto strike-slip faults (Figure 2.1b). The ST is bounded by the PRB to the west and by the Chocolate Mountains to the northeast (Crowell and Sylvester, 1979). Seismicity in the ST is abundant and is mainly above 14 km depth (Doser and Kanamori, 1986; Hauksson, 2011). Large earthquakes have been recorded on the main transform faults and frequent earthquake swarms occur in the small spreading centers south of the Salton Sea and likely indicate active diking (Hill, 1977). Subsurface magmatic activity is expressed at the surface as small Quaternary rhyolite volcanoes CHAPTER 2. BARAK ET AL., G-CUBED 2015 14 at the Salton Buttes (0.5–5.9 ka) (Wright et al., 2015) and Cerro Prieto (108 ka) (Damon and Shafiqullah, 1991). Analysis of xenoliths from the Salton Buttes and Cerro Prieto (Schmitt and Vazquez, 2006) supports the interpretation from seismic data that the underlying crust is composed of sediments and new oceanic crust (Fuis et al., 1984). The heat flow in the ST (average 140 mW/m2; locally ≥500 mW/m2) is significantly higher than in the adjacent ranges (Lachenbruch et al., 1985; Bonner et al., 2003). Hydrothermal circulation, evident at the surface as geothermal activity, may act to cool the crust and thus explain the only minor localized expressions of volcanism despite the high heat flow (Schmitt and Vazquez, 2006). Crustal and upper-mantle structure beneath the ST has been imaged by both active-source seismic studies (Biehler et al., 1964; Fuis et al., 1984; Parsons and Mc- Carthy, 1996; Rose et al., 2013; Han et al., 2013, (Figure 2.1b)) and passive-source seismic studies (e.g., Hauksson, 2000; Lin et al., 2007; Tape et al., 2009; Schmandt and Humphreys, 2010). The sedimentary basin in the Imperial Valley is 5–6 km deep with P-velocities (VP ) increasing from ∼2.0 to 5.2 km/s. The basement beneath the

Imperial Valley has a low average VP of 5.65 km/s, and is believed to be composed of a mix of metasedimentary rocks and minor amounts of igneous rocks. A deep layer with VP of 6.8–7.2 km/s inferred from seismic and gravity measurements has been interpreted as diabase above gabbro (Fuis et al., 1984). Lower-crustal VP of 7.2–7.4 km/s was also observed south of the border (Ramirez-Ramos et al., 2015).

Tape et al. (2009) found mid-crustal VS of >3.9 km/s beneath the ST basin, in ac- cord with the high VP , but their model has relatively low VS of ∼3.5 km/s in the deepest crust. Lin et al. (2007) and Allam and Ben-Zion (2012), also show high VP and high VP /VS from ∼10 to 17 km beneath the ST. The crustal thickness of the ST decreases from 27 km at the eastern margin beneath the Chocolate Mountains to ∼21 km (Parsons and McCarthy, 1996) or even as thin as 17–18 km (Han et al., 2013) beneath the Imperial Valley, to 25 km beneath the ePRB along the western edge of the ST (Ichinose et al., 1996). The upper-mantle beneath the Imperial Valley has low VP (7.6–7.7 km/s) (Parsons and McCarthy, 1996), or even lower (7.5 km/s) (Han et al., 2013). Schmandt and Humphreys (2010) observed prominent low VP and VS

(high VP /VS) anomalies in the mantle, extending to 150–200 km depth beneath and west of the Salton Trough and estimated ∼1% partial melt. Post-seismic relaxation CHAPTER 2. BARAK ET AL., G-CUBED 2015 15 of the M=7.2 2010 El Mayor-Cucapah earthquake implies the weakest upper mantle beneath the Imperial Valley has a viscosity 1.5 orders of magnitude lower than the strongest upper mantle to the west beneath the PRB and to the east beneath the Southern Colorado River Desert (Pollitz et al., 2012). This decrease in viscosity could correspond to a sub-solidus temperature increase of 100–200◦C, or a melt fraction of 10–20%, or a change from dry to wet conditions (e.g., B¨urgmannand Dresen, 2008).

2.4 Data and Method

In this section, we describe the ambient-noise data used in our analysis, followed by our data-processing procedures, tomographic inversion, and the conversion of surface- wave results obtained at diﬀerent periods to a 3-D model of shear-wave velocities.

2.4.1 Data

We use vertical-component data, archived at the IRIS DMC and SCEC Data Center, recorded at 849 broadband stations in and around southern California, spanning three different time periods: 1997–1998, 2007, and 2011 (Figure 2.1a). These specific time periods were chosen to augment available permanent stations with data from campaign broadband arrays in order to optimize ray-density and azimuthal coverage around the Salton Trough region, southern California. The first time period (1997– 1998) includes, among other stations, the Northern Baja Transect (Lewis et al., 2001), providing more data south of the U.S-Mexico border, and the LARSE II passive stations (Murphy et al., 2010), a dense broadband array to the north. The second period (2007) includes USArray stations for additional data and coverage to the north and the east. The last period (2011) includes our own broadband Salton Seismic Imaging Project (bbSSIP), a dense broadband array deployed across and around the Salton Trough.

2.4.2 Seismic Waveform Processing

For each of the time periods described above, we process the data to extract noise correlation functions (NCF) following Seats et al. (2012). We organize the seismograms into day-long records, remove the instrument response, and downsample the data to CHAPTER 2. BARAK ET AL., G-CUBED 2015 16

1 Hz. We window the seismograms into hour-long time series with a 75% overlap, de- mean them, and discard data windows that contain an amplitude spike greater than 8 times the standard deviation. We then calculate the coherency in the frequency domain for every station-pair and for each hour-long window. We limit our output to 2500 s (larger than twice our largest inter-station distance divided by the slowest estimated Rayleigh-wave group-velocity). Finally, we stack the coherences calculated for every station-pair to obtain an estimate of the earth’s response between them. We only use stacks that contain at least 40% of the original data windows and span at least 10 days. Because the distribution of noise sources, especially at short periods, can be very directional, especially in southern California where most of the noise is coming from the Pacific (e.g., Schulte-Pelkum et al., 2004), we average the causal and acausal sides of the NCF. Our averaging allows for automation of the process without the need to check both sides of the NCF for consistency. Although averaging tends to decrease the signal quality when the noise source is directional (Supp. figures S1), Pawlak et al. (2011) reports only a minor difference between averaging and choosing the side with the best signal quality. For each station pair, we calculate an average Rayleigh-wave group-velocity dispersion curve between the two stations using the multiple-filter technique (Dziewonski et al., 1969). In this process, we constrain the group-velocities to be within 1 and 5 km/s, to abate outliers. Our maximum allowed velocity of 5 km/s was chosen in part to mitigate any bias due to fast body wave arrivals (Hillers et al., 2013) (supporting information Figure S1). We also exclude results for which the inter-station distance is smaller than 3 wavelengths in order to ensure we satisfy a far-field approximation.

2.4.3 2-D Tomographic Inversion

We combine the inter-station dispersion curves from all three epochs to derive a group-velocity layer model using an iterative conjugate-gradient least-squares inversion (Paige and Saunders, 1982). Each period is inverted separately, using the infinite- frequency raypath approximation and using the perturbation from relative velocity from the mean. The data are weighted by the number of raypaths in each cell, normalized by the mean number of raypaths. Spatial smoothing is applied using the first-derivative roughness matrix (e.g., Lawrence and Prieto, 2011). We tested a range CHAPTER 2. BARAK ET AL., G-CUBED 2015 17 of model-smoothing parameters, and chose a smoothing parameter of 2 for our final model. We chose the smoothing parameter that corresponded to the least-smooth model that still contained reasonable velocities consistent with the range of previous ANT models (Moschetti et al., 2007; Shapiro et al., 2005; Sabra et al., 2005). The resulting group-velocity model spans the area of Figure 2.1a, with 0.05◦ (∼5 km) x 0.05◦ lateral block size, spanning from 4 to 40 s period, at 2 s increments. Because the inversion was unstable when the entire data set was used, we statistically excluded outlier dispersion curves. After some trial and error, we discarded dispersion curves that had a velocity jump larger than 0.3 km/s between adjacent periods (2 s interval), and at each period we discarded group-velocity measurements that fell outside twice the interquartile range of group velocities in that period. We also performed checkerboard resolution tests to examine velocity recovery for different period slices (supporting information Figures S4–S41).

2.4.4 Inversion to VS(z)

In the southern California area of our group-velocity model (Figure 2.1b), we estimated the 1D shear-velocity depth-structure at each grid cell, using surf96 from the CPS code package (Herrmann, 2013). An initial model of VP , VS and density is required to guide the 1D inversion process. We used CVM-H11.9, a high-resolution community model of southern California (Plesch et al., 2011), as an initial model. In order to mitigate the bias of our results to CVM-H11.9, we created a suite of 13 starting models based on CVM-H11.9. The suite of 13 models is comprised of the original model, 3 models with higher and 3 models with lower velocities and densities than CVM-H11.9 (by 5%, 10% and 15%), maintaining the same velocity gradient, 3 models with higher velocity gradients and lower surface velocities than CVM-H11.9, and 3 models with lower velocity gradients and higher surface velocities than CVM-H11.9 (Figure 2.2). We inverted for a model of 1 km thick layers down to 70 km depth from the surface. We also included six 10 km thick layers below 70 km, above a half-space, to avoid undesired eﬀects at the bottom of the model. In all the inversions we used the default damping parameter (1.0) for 100 iterations, except in the ﬁrst iteration for which we used a damping parameter of 100 in order to avoid overshooting (following

Herrmann (2013)). Our ﬁnal VS model, presented in this paper, is the median result CHAPTER 2. BARAK ET AL., G-CUBED 2015 18

0 Original Increase Decrease 10 Decrease −> Increase Increase −> Decrease

−5% +5%

30 −10% +10%

−15% +15% Depth [km] 40

345 Vs [km/s]

Figure 2.2: Schematic of 13 starting models used for each 1-D profile of VP , VS and ρ extracted from CVM-H11.9, created by increasing and decreasing the velocity up to 15%. CHAPTER 2. BARAK ET AL., G-CUBED 2015 19 of all 13 inversions, interpolated between 0 km depth (mean sea level) and 50 km below sea level. We also use the standard deviation of the 13 models as an estimate of the uncertainty we have at each point. Differences between the final model and CVM-H11.9 are considered significant if they are larger than the standard deviation of the 13 models. Since our 3D VS model is composed of individual 1D inversions, the horizontal resolution of the model is governed by the grid spacing (∼5 km), and by the lateral resolution of the group-velocity model. Therefore, the raypath density and checkerboard maps of the group-velocity model presented in the supporting information are a good proxy for horizontal resolution. Misfit histograms for most periods and a model resolution matrix for the average model are also available in supporting information.

2.5 Rayleigh Group-Velocity Model

Our group-velocity model is in agreement with other previous surface-wave models, but has better resolution. We compare the full model, encompassing most of the western U.S, with Moschetti et al. (2007) (Figure 2.3). First, we note that we used vastly more stations in California, Utah, Arizona, southern Oregon and to the east, because we included additional permanent stations (not part of the USArray Transportable Array), campaign stations and additional USArray stations that have propagated east since the Moschetti et al. (2007) analysis (Figure 2.1a). Also, our lateral grid spacing is 0.05◦, whereas their grid spacing was 0.5◦. The two models are broadly consistent, but we observe ﬁner details where we have better ray coverage, mostly in southern California. We have masked parts of the model for which the cumulative length of ray-paths within a cell is less than 0.1◦ (∼11 km) (Figure 2.4). We also compared the southern California part of our model with more local ambient-noise studies of southern California (Shapiro et al., 2005; Sabra et al., 2005; Zigone et al., 2015), compared to which we obtain higher resolution over a wider range of periods (supporting information Figures S4–S41). We note prominent features in our group-velocity model, relevant to the following interpretation of our ﬁnal shear-wave model (Section 2.6). Short periods are sensitive to shallow depths and long periods to deeper depths. First, we observe striking CHAPTER 2. BARAK ET AL., G-CUBED 2015 20

(a) Our Final Group−Velocity Model (b) Moschetti et al., 2007

44˚

40˚

36˚

32˚ 14 sec 14 sec Mean is 2.90 km/s Mean is 2.87 km/s

−124˚ −120˚ −116˚ −112˚ −108˚ −124˚ −120˚ −116˚ −112˚ −108˚ dU% −25 −20 −15 −10 −5 0510 15

Figure 2.3: (left) Comparison of our Rayleigh group-velocity model to (right) Moschetti et al. (2007) showing deviation from the mean group velocity (U) at 14 s. Dashed-line rectangle: area of the model we inverted for shear-wave velocity. Solid thin lines: state and physiographic boundaries after USGS (1946). CHAPTER 2. BARAK ET AL., G-CUBED 2015 21 velocity contrasts across much of the SAF out to 28 s period. At 16 s (Figure 2.4b), we observe a prominent and continuous velocity contrast along the SAF trace, from its southern termination at the southern edge of the Salton Sea to its intersection with the Garlock Fault, always with higher velocities to the SW and lower to the NE (Figure 2.4b). Our model also shows high velocities beneath the PRB at all periods. North of the United States/Mexico border, we also observe a prominent west-east velocity contrast, higher velocities to the west, coincident with the compositional boundary that divides the western and eastern sections of the batholith Silver and Chappell (1988). As expected, the SNB, the northern extension of the PRB, has faster velocities than the adjacent southern San Joaquin basin of the Great Valley (GV) at short periods (Figures 2.4a and 2.4b). Beneath the basin, velocities are faster at long periods than to the east (Figure 2.4d), analogous to the west-east contrast from faster wPRB to ePRB. Finally, beneath the ST we observe low velocities at short periods corresponding to the thick sequence of basin sediments and underlying metasediments. At the longest periods beneath the ST (Figure 2.4d), we observe a prominent low-velocity anomaly which most likely represents hot upper-mantle.

2.6 Shear-Wave Velocity Model

We next present our improved shear-wave velocity model of southern California down to 50 km, gridded at 0.05◦ (∼5 km) x 0.05◦ x 1km vertically. In many seismic imaging studies, the subsurface structure is not well known in advance, and a global 1D model, such as PREM (Dziewonski and Anderson, 1981), or a crude 3D model, such as crust 1.0 (Laske et al., 2013), is used as a starting model for the inversion. We, however, used CVM-H11.9, a high-resolution (upper-crust: 1km x 1km x 100m vertical; lower crust/mantle: 10km x 10km x 1km vertical) 3D model of southern California as basis for our 1D inversion from Rayleigh-wave group-velocity to shear-wave velocity. Since we began this work, a new community model CVM-S4.26 (Lee et al., 2014) has been published, and in supporting information to this paper we compare our model to theirs. Our ambient-noise surface-wave data have limited vertical resolution (supporting information Figures S3). Therefore small-scale vertical velocity variations and sharp CHAPTER 2. BARAK ET AL., G-CUBED 2015 22

−120˚ (a) Upper Crust −116˚ −120˚ (b) Mid Crust −116˚ 36˚ 36˚ S S N N B B GV GV

GF GF

SAF SAF

ST ST PR PR

B B IV IV

32˚ 08 sec 16 sec 32˚ dU% dU% −20 −10 010 −10 010 (c) Lower Crust (d) Upper Mantle 36˚ 36˚ S S N N B B GV GV

GF GF

SAF SAF

ST ST PR PR

B B IV IV

32˚ 24 sec 40 sec 32˚ dU% dU% 01020 −2 0 246 Figure 2.4: Maps of our Rayleigh group-velocity model in southern California, at periods of 8, 16, 24, and 40 s, shown as deviation from the mean group velocity (U) of the entire western U.S. model for that period, with color range varying between maps due to the large variation in velocity perturbations. Annotations as in Figure 2.1b. CHAPTER 2. BARAK ET AL., G-CUBED 2015 23 velocity contrasts (e.g., Moho), that are present in the CVM-H11.9 starting model, appear in our final model largely unchanged (we compare the models in the supporting information figures). Although many other geologically significant features are undoubtedly present in our model, we only discuss features that (1) are laterally extensive, ∼50 km or 10 grid points wide, (2) are observed in our intermediate group-velocity model (Figure 2.4), for which our checkerboard tests (supporting information Figures S4–S41) confirm we have good model resolution at 0.5◦ or better, (3) show a difference from CVM-H11.9 that is greater than the standard deviation of our suite of 13 inversions of our 13 starting models (Figure 2.2) (deviations are shown in supporting information Figures S42–S58), and (4) have, we believe, ready geological interpretations. CHAPTER 2. BARAK ET AL., G-CUBED 2015 24

−120˚ −118˚ −116˚ −114˚ Figure 2.5: Depth slices through our ﬁnal 36˚ Vs model at 7, 20 and 42 km depth.

(a) 7 km 32˚ −120˚ −118˚ −116˚ −114˚ 36˚

(b) 20 km 32˚ −120˚ −118˚ −116˚ −114˚ 36˚

Section 2.6.1 presents interpretations of 2D profiles AA0 through II0 across the SAF (Figure 2.1c). Section 2.6.2 presents interpretations of 2D profiles PRaa0–PRcc0 across the PRB, and 2D profiles SNaa0 and SNbb0 across the SNB. Section 2.6.3 discusses profiles STaa0–STcc0 across the ST.

2.6.1 San Andreas Fault

Model Description

We observe changes in velocity structure along the southern SAF that lead us to divide the fault into three distinct sections (Figure 2.1c): southeast of San Gorgonio Pass (Figures 2.6a– 2.6c), along the San Bernardino Mountains (Figures 2.6d– 2.6f), and northwest of the San Bernardino Mountains (Figures 2.6g– 2.6i). Although we only show a few exemplary cross-sections, we have examined 2D profiles perpendicular to the SAF every 10 km along the entire length of the southern SAF. Our southern section of the southern SAF (profiles AA0, BB0, CC0), is character- ized by striking velocity contrasts (∼3%) across the SAF, with vertical-to-steep NE dips (Figures 2.6a– 2.6c and 2.7a). We observe the large lateral velocity gradients throughout the crust and in the upper mantle (e.g., Figure 2.6b). Beneath the deepest seismicity (ca. 15 km), our velocity contours become gradually separated with depth (corresponding to decreasing lateral velocity gradient, dv/dx), and the velocity contours change from steeper to shallower. In comparison, the CVM-H11.9 model shows only negligible lateral velocity gradients in the lower crust and none in the mantle. Along this section of the SAF, our observed velocity contrast has the same general location and dip as that predicted from the Fuis et al. (2012) SAF model (Figure 2.6, thick-dashed line). In our central section of the southern SAF (profiles DD0, EE0, FF0), we observe an unusually high-velocity zone beneath the bottom of seismicity that dips ∼20◦ NE. This high-velocity structure seems to be part of the high-velocity lower crust observed beneath the PRB (Figure 2.7b) (See section 2.6.2). Our velocity contrasts along this section do not align well with the Fuis et al. (2012) SAF model, although both their model and our isovelocity contours have a NE dip. We note that the NE-dipping boundary between higher (blue) and lower (green) velocities in the lower crust appears to extend into the upper mantle. CHAPTER 2. BARAK ET AL., G-CUBED 2015 26

In our northernmost section of the southern SAF (profiles GG0, HH0,II0), the velocity contrast across the fault becomes less obvious, and is not consistent along the fault. Where the SAF flanks the San Gabriel Mountains to the NE there may be a vertical velocity contrast aligned with the Fuis et al. (2012) model (Figure 2.6h). Further to the NW, where the Fuis et al. (2012) model dips SW, we observe some poorly defined SW-dipping iso-velocity contours aligned with the SAF model from Fuis et al. (2012) (Figure 2.6i). Although not as clear as in the central section of the SAF, a boundary between high and low velocities in the mantle is also seen along the projection of the model SAF of Fuis et al. (2012).

Interpretation and Discussion

We ﬁrst emphasize that the lateral velocity contrasts across the SAF in our ﬁnal shear-wave model (Figure 2.6) seem to be largely driven by the data, as we see them in our group-velocity model (Figure 2.4), but they are largely absent from the CVM- H11.9 model. In addition, the along-strike consistency and gradual variation of our velocity structure indicate that these features are not an artifact of a single bad 1D inversion. Along our southern section of the southern SAF (Figures 2.6a– 2.6c), we observe large lateral velocity gradients throughout the crust and in the upper mantle. We infer that this velocity contrast represents the SAF reaching the Moho or deeper, as a relatively narrow structure. Our observed velocity contrast is in general agreement with the Fuis et al. (2012) SAF model, but our continuous 3-D model may help

Figure 2.6: True-scale cross-sections perpendicular to the San Andreas Fault along lines shown in Figure 2.1c. (left) ﬁnal model. (right) CVM-H11.9 model, the basis for our starting models in our shear-wave inversion. Iso-velocity contours are at 0.05 km/s intervals. Magenta circles are seismicity from Hauksson et al. (2012), aggregated over a width of 5 km and scaled by magnitude. Cyan lines are faults from CFMv5 (Plesch et al., 2007). Heavy dashed-line shows the dip and depth extent of the SAF from Fuis et al. (2012). Above each cross section we show topography (greatly vertically exaggerated) in shaded gray (mean sea level: thin-dotted line), Bouguer gravity (solid line) (Pan- American, 2010), and magnetic potential (pseudogravity, dashed line) (Langenheim et al., 2014). Supporting information Figures S42–S50 provide standard deviation of our model and comparison with CVM-S4.26 (Lee et al., 2014). Abbreviations as in Figure 2.1b. Additional faults not annotated in Figure 2.1b: BRF: Buck Ridge Fault; CF: Clark Fault; CFN: Clark Fault North; CCF: Coyote Creek Fault; CTF: Claremont Fault; JVF: Johnson Valley Fault; MHF: Mecca Hills- Hidden Springs Fault; SGPD: San Gorgonio Pass Detachment; SSGF: Southern San Gabriel Fault; WF: Whittier fault. CHAPTER 2. BARAK ET AL., G-CUBED 2015 27

[pmGal] A Final A' A CVM−H11.9 A' SW Bouguer Mag. Pot. NE

10 −40 [mGal] 5 −50 Potential 0 V.E.=15 −60

EF SAF EF SAF Bouguer 0 0 Magnetic 3.5 3.5 10 3.5 10 3.65 3.5 3.6 (a) 3.7 20 3.7 20 4.35 3.55 km 4.35 4.5 4.5 4.5 km 30 30 4.45 4.35 4.2 4.4 40 4.1 4.3 40 4.25 4.35 4.3 50 V.E.=1 V.E.=1 50 020406080 100 020406080 100 km km

[pmGal] B B' B B' SW NE −40 10 [mGal] −60 5 Potential V.E.=14 −80

CCF CF SAF MHF CCF CF SAF MHF Bouguer 0 0

Magnetic 3.5 3.5 3.5 10 10 3.65 3.5 (b) 3.65 20 20 3.7 3.7 4.45 3.5 3.5 3.65 4.5 3.65 km 4.4 5 km .3 4. 4.5 4.5 30 4.4 4 30 4.35 4. 40 4.15 4.2 4.35 4 40 4.1 4.3 4.35 50 V.E.=1 V.E.=1 50 020406080 100 020406080 100 km km

[pmGal] C C' C C' SW NE 30 −70 [mGal] V.E.=6 Potential 10 −90

CCF CF BRF SAF MHF CCF CF BRF SAF MHF Bouguer 0 0 3.5 3.5 Magnetic 10 3.65 10 3.5 (c) 3.5 3.65 20 3.7 3.6 3.65 20 3.6

km 3.7 3.8 3.7 km 4.3 3.7 4.35 4.45 30 4.45 30 4. 4.25 4.5 2 4.4 4.25 4.4 5 40 4.15 40 4.1 4.25 50 V.E.=1 V.E.=1 50 020406080 100 020406080 100 km km

Vs [km/s]

0.5 1.0 1.5 2.0 2.5 3.0 3.5 4.0 4.5 Figure 2.6: Caption on previous page. CHAPTER 2. BARAK ET AL., G-CUBED 2015 28

Final CVM−H11.9 D D' D D' SW NE Bouguer Mag. Pot. [mGal]

Potential 30 −80 −100 V.E.=8

[pmGal] 10

CFN SAF JVF CFN SAF JVF Bouguer Magnetic 0 0 3.5 3.5 10 3.5 3.5 10 3.7 (d) 3.65 20 3.8 SGPD 3.6 SGPD 20

km 3.8 3.6 km 3.75 3.7 30 3.8 30 3.9 3.95 4.25 4.35 40 4.45 40 4.4 50 V.E.=1 V.E.=1 50 020406080 100 020406080 100 km km E E' E E' SW NE [mGal] Potential 30 −60 20 V.E.=6 [pmGal] −100 10

CTF SAF CTF SAF Bouguer Magnetic 0 0 3.5 10 3.5 10

(e) 3.55 20 SGPD 3.65 SGPD 20 3.85 3.7 km km 3.75 30 3.95 3.75 30 3.95 3.8 4.2 40 4.45 4.25 40

4.3 4 50 V.E.=1 V.E.=1 50 020406080 100 020406080 100 km km F F' F F' SW NE Potential 30 [mGal] −80 20 [pmGal] 10 V.E.=10 −120 Bouguer

Magnetic EF SJF SAF EF SJF SAF 0 0 3.5 3.5 10 3.7 3.5 10 3. 3.6 5 (f) 20 3.55 20 km km 3.7 3.7 3.9 3.85 30 5 30 3.9 4.05 3.8 40 4.35 40 4.45 4.2 4.25 4.4 50 V.E.=1 V.E.=14.4 50 020406080 100 020406080 100 km km

Vs [km/s] 0.5 1.0 1.5 2.0 2.5 3.0 3.5 4.0 4.5 Figure 2.6: CHAPTER 2. BARAK ET AL., G-CUBED 2015 29

G Final G' G CVM−H11.9 G' SW Bouguer Mag. Pot. NE

10 [mGal]

Potential −80 −100

−10 V.E.=6 [pmGal]

WF SAF WF SAF Bouguer Magnetic 0 0 3.5 3.5 10 3.5 3.5 3.5 10

(g) 7 20 3. 20

km 3.65 3.8 km 30 3.6 3.8 30 3.65 3.7 3.75 4 4.2 .3 40 4.35 4.4 40 4.2 4.4 50 V.E.=1 V.E.=1 50 020406km 080 100 020406km 080 100 H H' H H' SW NE [mGal] Potential −60 −5 −100 [pmGal] −15 V.E.=10

SSGF SAF SSGF SAF Bouguer Magnetic 0 0 3.5 3.5 10 10 (h) 3.65 3.65 3.5 3.5 20 3.7 20 3.7 km km 3.6 3.55 3.65 3.7 3.7 30 4.5 3.8 30 4.3 4.35 4.45 40 4.2 4.4 40 4.3 50 V.E.=1 V.E.=1 50 020406km 080 100 020406km 080 100 I I' I I' SW NE

−60 [mGal]

V.E.=6 −100

SAF SAF Bouguer 0 0 3.5 3.5 10 3.5 10 (i) 3.55 3.5 20 20 3.7 3.7 3.8 km 3.6 km 3.8 .7 30 4.45 3 30 4.45 3.75 4.3 4.55 4.3 40 5 40 4 4. 4.35 50 V.E.=1 V.E.=1 50 020406km 080 100 020406km 080 100

Vs [km/s] 0.5 1.0 1.5 2.0 2.5 3.0 3.5 4.0 4.5

Figure 2.6: CHAPTER 2. BARAK ET AL., G-CUBED 2015 30 reﬁne the Fuis et al. (2012) model, which consists of 15 discrete 2D cross-sections along a 300-km section of the fault. The velocity contrast across proﬁle BB0, for example, suggests a steeper fault than Fuis et al. (2012) and also aligns better with the projected seismicity pattern, thereby implying a more rapid change in dip along the fault than previously suggested. Our results also support the projection of the SAF into the upper mantle previously inferred by Fuis et al. (2012). The decrease of lateral velocity gradient with depth that we observe may suggest a shear zone of increasing width. However, the resolution of our starting model (CVM-H11.9) decreases in the lower crust. It is therefore possible that the reduction in both lateral and vertical velocity gradient in our model is an inherited change in resolution. The steeply NE-dipping velocity contrast that we observe is however also very evident in the recent, rougher model CVM-S4.26 (Lee et al., 2014) (supporting information Figures S42–S44). Along our middle section of the SAF, we observe a high-velocity lower-crustal wedge. The top of this wedge structure is aligned with the base of seismicity where a lower-crustal detachment (the “San Gorgonio detachment”) has been inferred, e.g., (Seeber and Armbruster, 1995) (Figure 2.6e). This high-velocity lower-crustal wedge is likely underthrust western Peninsular Ranges lower-crust, or possibly a remnant of the Farallon plate (see Section 2.6.2). NE-underthrusting is consistent with the direction of lower-crustal anisotropy inferred southwest of the San Andreas Fault (Porter et al., 2011). The NE-directed underthrusting along the inferred detachment may have helped rotate the SAF into its local NE dip. We note that our observed high-velocity structure continues to dip NE into the upper mantle, a behavior not present in our starting model (Figures 2.6e– 2.6f), even though the Moho depth in our model is essentially locked to the Moho from the initial model. This feature appears to be consistent with the observations of Fuis et al. (2012) whereby the modeled SAF in the crust appears to extend into the mantle to the NE side of the high-velocity mantle body of Kohler et al. (2003). In our northern section (Figure 2.6g– 2.6i), velocity contrasts across the SAF are smaller and more laterally variable along the fault trace than further south. This behavior is consistent with the inferences of Fuis et al. (2012) that velocity contrasts are small across the SAF (≤0.1 km/s) and that the fault dip varies from SW-dipping CHAPTER 2. BARAK ET AL., G-CUBED 2015 31 to NE-dipping over a distance of less than 100 km.

2.6.2 The Mesozoic Batholiths: Peninsular Ranges and Sierra Nevada

Model Description

S-velocities beneath the PRB and SNB are higher than anywhere else in our study area, as observed in both our shear-wave model (Figure 2.5) and group-velocity model (Figure 2.4). We observe a prominent west-east velocity variation beneath both the

PRB and SNB, vertically spanning most of the crust and upper mantle. VS reaches 4.05 km/s in the lower crust and 4.6 km/s in the upper mantle beneath the wPRB, while beneath the ePRB VS remains low (≤3.7 km/s) throughout the crust and reaches only 4.45 km/s in the upper mantle. The velocity transition beneath the PRB clearly dips to the east and projects to the surface along a known compositional boundary (e.g., Silver and Chappell, 1988; Langenheim et al., 2014) (Figures 2.7 and 2.8). At the northeastern edge of the wPRB, the high-velocity zone dips more gradually (∼20◦) and continues northeast, beneath the SAF, and under the San Bernardino Mountains (Figures 2.6d– 2.6f, 2.7b and 2.8b). The southern termination of the high-velocity zone beneath the wPRB (Figure 2.5) is apparent in our group-velocity model at periods ≥16 s (Figures 2.4c and 2.4d). Similarly, beneath the SNB, we observe higher crustal and upper-mantle S-velocities (up to 4 km/s in the lower crust) beneath the wSNB and eastern Great Valley relative to the eSNB (Figure 2.10b). As in the PRB, the west-east velocity transition seems to dip to the east. In the upper mantle, velocities are highest beneath the eastern Great Valley. Our west-east proﬁle (Figure 2.10b) beneath the SNB and eastern Great Valley lies along proﬁle 6 of Fliedner et al. (2000) (Figure 2.10a). Although we do not image the upper crustal high-velocity zone observed in Fliedner’s model beneath the Great Valley and interpreted as Great Valley ophiolite, their model shows higher velocities beneath the wSNB relative to the eSNB along an east dipping boundary in agreement with our model. High S-velocities are also evident in the lower crust beneath the GV in CVM-S4.26 (supporting information Figures S54–S55). CHAPTER 2. BARAK ET AL., G-CUBED 2015 32

36˚ SNb SNa' SNa (a) SNb'

34˚ STa PRa PRa' STc STb'

PRb PRb’

PRc PRc' STb 3.6 km/s STa 32˚ STc' ' −120˚ −118˚ −116˚ −114˚

10 Depth [km] 20 30

36˚ SNb SNa’ SNa

(b) SNb’

34˚ STa PRa PRa’ STc STb’ 25 PRb PRb’

PRc PRc’ STb STa’20 3.8 km/s STc’ 32˚ −120˚ −118˚ −116˚ −114˚

10 20 30 Depth [km]

Figure 2.7: (a) Depth to the shallowest 3.6 km/s isovelocity surface, showing deepening (lower velocities) NE of the SAF and NE of the boundary between wPRB and ePRB. (b) Depth to the shallowest 3.8 km/s isovelocity surface, showing deepening (lower velocity) NE of the PRB and E of the prebatholithic suture in the SNB (see Figure 2.10). Fault traces and physiographic divisions as in Figure 2.1. Depth contours at 5km interval. CHAPTER 2. BARAK ET AL., G-CUBED 2015 33

PRa Final PRa'

W Bouguer Mag. Pot. E 30 −20 10 −60 [pmGal] −10 V.E.=6

EF SJF SAF Bouguer [mGal] Magnetic Potential wPRB ePRB 0 0 3.5 3.5 10 3.5 10 3.55 3.85 3.75 20 3.95 3.8 20

(a) km 3.85 3.8 km 30 4.1 3.9 4.4 30 4.45

40 4.55 40 4.3 4.15 4.35 50 V.E.=1 50 020406080 100 120 140 160 km PRb PRb' Vs [km/s] 4.5 W E −20 50 −40 0 V.E.=6 −60 [pmGal] 4.0 coast wPRB EF ePRB SJF Bouguer [mGal] Magnetic Potential 0 0 3.5 3.5 3.5 10 10 3.5 3.7 3.8 20 3.6 20 3.95 km 4.45 km (b) 30 3.85 30 3.0 3.95 4.05 4.25 2 40 4.6 4. 4.1 40

4.55 2.5 50 V.E.=1 50 02040608km 0 100 120 140 PRc PRc' 2.0 100 −20 50 −60

[pmGal] V.E.=8 0

LSF EF Bouguer [mGal] Magnetic Potential W wPRB ePRB E 1.5 0 0 3.5

10 10 3.5 1.0 20 20

km 3.75 3.5 km (c) 3.8 30 3.9 4.35 30 3.9 0.5 4.4 40 4.35 40 4.45 4.2 50 V.E.=1 50 02040608km 0 100 120 140

Figure 2.8: West-east proﬁles across the PRB through our VS ﬁnal model, as shown in Figure 2.1c. Figure annotation as in Figure 2.6. Dashed line: compositional boundary between wPRB and ePRB, extrapolated to depth at a 45◦ angle (cf. Langenheim et al., 2014). CHAPTER 2. BARAK ET AL., G-CUBED 2015 34

PRb Bouguer PRb' −20 50 Mag. Pot. −40 −60 [pmGal] 0 Bouguer [mGal] Magnetic Potential coast EF SJF 0 3.25 3 0 Vs [km/s] 3 3 3.75 10 10

20 20 3. 4 km 3.75 km 9 30 30

(a) Final 3.95 40 4.55 4.1 40

44 50 V.E.=15 50 020406080 100 120 140 3 0 3 0 3 3.25 3 10 10 3. 20 7 20 3.65 30 3.8 30 3.75 2 40 4.35 40

(b) CVM−H11.9 4.35

50 V.E.=1 50 020406080 100 120 140

0 3.5 3 0 3 3.45 1 10 10 3.5 20 3.65 20

30 4.0 3.7 30 5 3 40 4. 40 (c) CVM−S4.26 4.5 50 V.E.=1 4.5 50 020406080 100 120 140 S.D. [%] 0 0

10 10 4

20 20 km km 30 30 2

40 40 (d) S.D. our model 50 V.E.=1 50 0 020406080 100 120 140 km Figure 2.9: Comparison along PRb–PRb0 (Figure 2.8b) between our model (a), CVM-H11.9 (initial model; b), and CVM-S4.26 (c). (d) Standard deviation of our model. Figure annotation as in Figure 2.8. See also supporting information Figures S51–S53. CHAPTER 2. BARAK ET AL., G-CUBED 2015 35

Discussion: Peninsular Ranges

The velocity structure of our model beneath the PRB contrasts greatly with the CVM- H11.9 model (Tape et al., 2009), the initial model we used for the shear-wave inversion, and with the CVM-S4.26 model (Lee et al., 2014) (Figure 2.9 and supporting information Figures S51–S53). The CVM-H11.9 model has low velocities in the lower crust beneath higher velocities in the mid-crust, while in our model the velocities generally increase with depth beneath the PRB (Figure 2.9b). Because CVM-H11.9 is poorly constrained below the seismogenic zone, we suspect their lower-crustal velocities are an artifact of too much horizontal smoothing. The CVM-S4.26 model, although more heterogeneous than our model, shows much higher velocities than our model (Fig- ure 2.9c), which we suspect are due to too little smoothing. The S-velocities that reach 4.3 km/s above 20 km depth in CVM-S4.26 would imply ultramaﬁc (mantle) rocks, whereas our S-velocities of ∼3.8–3.9 km/s are in better agreement with the VP of 7 km/s lower-crustal velocity of Nava and Brune (1982) (Figure 2.1b), assuming

VP /VS of about 1.8, appropriate for mafic rocks (Christensen, 1996). In addition, the tomographic model of Zigone et al. (2015) shows velocities in the shallow crust of the PRB very similar to ours, and does not support the extraordinarily high velocities (> 3.9 km/s) shown in CVM-H11.9 and CVM-S4.26. Our velocity model corroborates potential-field models that indicate that the west- east compositional boundary in the PRB dips to the east and extends to the base of the crust (Figure 2.8) as previously speculated (Langenheim et al., 2014; Magistrale and Sanders, 1995). However, if the boundary between wPRB and ePRB is parallel to our iso-velocity contours, then the boundary dips at a shallower angle than the ∼45◦ previously inferred from gravity and magnetic modeling (Langenheim et al., 2014). Potential-field models likely have less resolution than seismic models in the lower crust. The lower-crustal, high S-velocities beneath the wPRB in our velocity model require mafic rocks (Christensen, 1996). The ∼15 km thick, high-velocity lower crust beneath the wPRB in our model (Figure 2.8) is similar to that observed in the oceanic

Aleutian arc, which has a ∼10 km thick lower crust with VS of 3.8–4.0 km/s (Flied- ner and Klemperer, 1999). Geobarometry in the wPRB (Ague and Brimhall, 1988) indicates that ∼4–11 km of crust has been removed, suggesting that if the Moho CHAPTER 2. BARAK ET AL., G-CUBED 2015 36 represents the crust-mantle boundary then the total thickness of arc crust was 35-40 km. Because the exposed wPRB plutons average a tonalitic composition (Silver and Chappell, 1988), and about half the modern crustal thickness is a significantly more mafic lower-crust, this marginal arc must have an overall composition more mafic then average continental crust (cf. Calvert et al., 2008). Hence our results appear consistent with formation of the wPRB on pre-existing oceanic crust at the continental margin (e.g., Johnson et al., 1999). In contrast, the ePRB has had ∼11–23 km of overlying material removed (Ague and Brimhall, 1988), and has S-velocities of 3.7 km/s all the way to the Moho (Figure 2.8), implying an original felsic crust of ∼35–40 km. The similar initial thickness we infer for the wPRB and ePRB suggest that the modern difference in Moho depth is largely a function of the greater erosional depth of the ePRB, driven by isostasy. The high topography and low gravity (Figure 2.8) over the ePRB must in part be due to the unusually felsic composition. The observed velocity structure of the wPRB is clearly a poor match to the ePRB even though both originated within the same arc system. This observation is consistent with very different magma-generating processes as well as perhaps different initial basements as inferred by Silver and Chappell (1988) and Grove et al. (2008).

The lower-crustal high VS beneath the wPRB seems to continue into the upper mantle (Figures 2.8a, 2.8b, 2.6e and 2.6f) presumably linking to the contrast in VP (also high to the west, low to the east) previously tracked from ∼60–150 km depth (Kohler, 1999; Kohler et al., 2003; Fuis et al., 2012). Although Kohler (1999), among others, suggested that the high-velocity body in the mantle is a downwelling associated with oblique convergence across the SAF, an origin of this body as remnant Farallon slab is also possible, as discussed in section 2.6.2. The southern decrease of high velocities beneath the wPRB observed in both our S-velocity and group-velocity models may partially be due to less-good station coverage south of the border (Figure 2.1a). However, gravity and magnetic values decrease south of San Diego to near the US/Mexico border, before they increase again (Langenheim et al., 2014). This suggests that the wPRB may be disrupted in this region, but not terminated (Langenheim et al., 2014). CHAPTER 2. BARAK ET AL., G-CUBED 2015 37

Discussion: Southern Sierra Nevada

The high-velocity (“Isabella”) anomaly in the upper mantle, imaged tomographically with increasing resolution for more than 20 years (e.g. Biasi and Humphreys, 1992; Jones et al., 2014) has been interpreted either as a remnant of Farallon oceanic lithosphere (Wang et al., 2013b) or as maﬁc (“arclogitic”) root extending beneath both the wSNB and eSNB at end Cretaceous time (Saleeby et al., 2012, 2013). Recent thermomechanical models of mantle-lithosphere removal to form the Isabella anomaly require, as a critical prediction, strength drops in the lower crust to facilitate geologically rapid delamination (Saleeby et al., 2012). In particular, the prediction by Saleeby et al. (2012) of a 10 km thick, felsic low-viscosity lower-crustal channel extending continuously beneath, and beyond the margins of, the southern SNB seems diﬃcult to reconcile with the velocities that increase east to west from 3.5 to 3.9 km/s

(VS, Figure 2.10b) and 6.4 to 7.2 km/s (VP , Figure 2.10a). Zones of partial melt in the lower-crust suggested by Saleeby et al. (2013) are also hard to reconcile with our observed velocities under the southwestern SNB (Figure 2.10b) that are character- istic of dry, unmelted gabbro (see section 2.6.3, Figure 2.13). Instead, we note the dimensions of the crustal high-velocity zone beneath the wSNB (100 km west-east, 20 km top to bottom) are commensurate with the high-velocity body we map beneath the wPRB. We argue that both these high-velocity bodies correspond to the maﬁc lower-crust of a continental-margin arc, or possibly to underthrust fragments of oceanic lithosphere, as also imaged and interpreted by Wang et al. (2013b). Along our north-south cross-section (Figure 2.10c), which corresponds to cross- section F6 of Saleeby (2003), Saleeby interpreted that the Rand schist is underthrust

Figure 2.10: VS profiles through our final model beneath the SNB. (a) West–east line P6 of Fliedner et al. (2000). Patterning of wSNB and eSNB, and prebatholithic suture (PBS, solid magenta) from Saleeby (2003). Heavy-dashed line is our interpretation of east-dipping PBS, as in (b). Figure 2.10b shows profile SNaa0 (Figure 2.1c) along line P6 of Fliedner et al. (2000) in Figure 2.10a. Pre- batholithic suture (PBS, solid magenta) as in Figure 2.10a and “low-viscosity channel” from Saleeby et al. (2012). Heavy-dashed line (PBS?) as in Figure 2.10a and patterning of SNB follows our interpretation. (c) Profile along SNbb0, corresponding to north-south line F6 of Saleeby (2003), in which ∼100 km of sinistral motion on the Garlock Fault was restored. Rand Thrust and Moho (magenta) from Saleeby (2003). Dashed line is schematic of our reinterpretation of the Rand Thrust, and patterning of SNB follows our interpretation in Figure 2.10b. Standard deviation of our model, and comparison with CVM-H11.9 and CVM-S4.26 shown in supporting information Figures S54– S55. All other annotations as in Figure 2.6, with abbreviations from Figure 2.1b. CHAPTER 2. BARAK ET AL., G-CUBED 2015 38

(a) Fliedner et al., 2000 (Vp, Line P6)

wSNB PBS KCF eSNB Great Valley OVF 0 0 6.6 10 6.0 10 7 20 6.8 6.2 20 km 6. 6.8 4 30 6.4 30 7.8 6.8 40 40 7.8 50 V.E.=1 50 0 20 40 60 80 100 120 140 160 180 200 km SNa (b) Final Vs (west-east) SNa' −50 −100 V.E.=5 −150

Bouguer [mGal] W Great Valley wSNB PBS SNbb' KCF eSNB OVF E 0

10 3.6 3.5 3.7 PBS? 20 3.6 3.9 3.8 km mafic lower crust 30 3.8 ? postulated channel low-viscosity 40 Isabella anomaly: 4.1 slab or drip? 4.3 4.3 V.E.=1 50 020406080 100 120 140 160 180 200 km Vp [km/s] Vs (c) Final Vs (north-south) 4.5

SNb SNb' 7 4.0 −80 −110 −120 V.E.=8 −120 6 3.5

GF Bouguer [mGal] N SNaa' S 3.0 5 0 0 eSNB? 3.5 2.5 10 10 4 Rand Thrust PBS? 3.5 2.0 20 20 wSNB? 3.6 3

3.8 km ? Moho 1.5 30 3.8 30 2 1.0 40 4 40 4.3 4. 4.3 4.3 1 0.5 50 V.E.=1 50 02040608km 0 100 01020

Figure 2.10: Caption on previous page. CHAPTER 2. BARAK ET AL., G-CUBED 2015 39

100 km northward from the outcrop of the Rand Thrust, beneath the SNB, and extends to the Moho. Although other models show a north-dipping velocity increase from the Garlock fault (CVM-H11.9) or velocity inversion (CVM-S4.26) (supporting information Figure S55), we do not image such a structure. Furthermore, our lower- crustal S-velocities (≥3.7 km/s) are too high to be Rand schist, based on laboratory measurements at elevated pressure (Godfrey et al., 2000), and instead indicate more maﬁc rock types (e.g., Christensen, 1996). If the lower-crustal high velocities at the north end of SNbb0 are part of the western SNB arc, as preferred by us and also interpreted by Saleeby (2003), then they belong structurally above the Rand Thrust that must dip more steeply northwards than suggested by Saleeby (2003), to reach the Moho within ∼60 km of the Garlock Fault, as shown in Figure 2.10c.

2.6.3 Salton Trough

Model Description

Within ST we observe low S-velocities (∼1–3 km/s) to as much as ∼5 km in the center of the basin. At the southeastern edge of the Salton Sea, beneath the Salton Buttes, a group of Quaternary volcanoes, we observe a shallow (∼7 km) high-velocity (up to ∼3.7 km/s) zone, ∼15km wide and ∼2 km thick, surrounded by lower velocities (Figures 2.5a and 2.11c). We observe similar high velocities at this depth south of the United States/Mexico border, beneath the Cerro Prieto volcano and adjacent areas (Figure 2.5a). Beneath the Brawley Seismic Zone (BSZ) and to the SE, we observe high mid-crustal velocities of ∼3.6–3.8 km/s (Figure 2.11a), though both

SW and NE of the basin axis (parallel to STa-a‘) VS remains <3.6 km/s to the base of the crust (Figure 2.11b). Where we observe VS ≥3.7 km/s in the midcrust, we also observe a low-velocity zone (∼3.5 km/s) in the lower crust below ∼15 km depth (Figure 2.5b, 2.11b, and 2.11c). Similar low velocities in the lower crust also appear in CVM-S4.26 at this location (supporting information Figure S76). In the upper mantle beneath the ST, and beneath a thin mantle “lid” that barely reaches a velocity of 4.5 km/s, we observe two spatially separated low-velocity zones

(Figure 2.11 and 2.12) with VS locally ≤4.05 km/s at ∼42 km depth both beneath the Salton Buttes, and beneath the Laguna Salada Basin, ∼30 km northwest of the Cerro Prieto volcano and geothermal ﬁeld. The northern anomaly beneath the CHAPTER 2. BARAK ET AL., G-CUBED 2015 40

STa Final STa' Mag. Pot. Bouguer 5 −40 V.E.=137

[pmGal] −5 −80

SB US/MEX Bouguer [mGal]

Magnetic Potential NW BSZ SE 0 0 3.55 3.5 3.55 3.5 10 3.75 10 3.7 3.5 3.5 20 3.6 3.85 20 4.45 3.7

km 4.5 km 4.45 4.45 (a) 30 4.4 30 4.25 4.3 40 4.15 4.2 40 4.05 50 V.E.=1 50 020406080 100 120 140 160 180 200 km STb STb' Vs [km/s] 15 4.5 5 −50 V.E.=12 Chocolate Mts. [pmGal] −5 −70

US/MEX LSF EF SJF BSZ SHF Bouguer [mGal] Magnetic Potential SW NE 4.0 0 0 3.5 10 3.5 10 3.5 3.6 20 3.5 20 3.5 (b) km 4.3 km 30 4.35 30 3.0 40 4.25 4.2 40 4.2 4.15 4.4 50 V.E.=1 50 2.5 020406080 100 120 140 160 km STc STc' 2 −30 2.0 V.E.=15 [pmGal] −2 LSB −50

SHF SB IF SJF US/MEX Bouguer [mGal] Magnetic Potential N S 0 0 1.5

3.5 10 3.65 3.5 10

1.0 20 3.5 3.7 20 3.65 (c) km 4.35 km 30 4.25 30

0.5 40 40 4 4.05 42 50 V.E.=1 4.3 5 50 020406080 100 120 140 160 km

0 Figure 2.11: VS profiles through our final model in the Salton Trough region. (a) STa–STa along the axis of the trough. (b) STb–STb0 perpendicular to the axis of the Salton Trough. (c) STc– STc0 north-south profile that show both low-velocity zones in the upper mantle described in the text. Vertical dashed lines: intersection of profiles. Note very different vertical exaggeration of part (Figure 2.11a). Standard deviation of our model, and comparison with CVM-H11.9 and CVM-S4.26 shown in supporting information Figures S56–S58. Abbreviations from Figure 2.1b. CHAPTER 2. BARAK ET AL., G-CUBED 2015 41

Figure 2.12: Horizontal slice at 42 km depth showing low-velocity zones in the upper mantle beneath the Salton Trough. Green lines: coast and state borders. Purple lines: surface trace of faults as in Figure 2.1b. Dot-patterned boxes: active spreading centers after Elders et al. (1972). Other annotations as in Figure 2.1b.

Salton Sea basin (Figure 2.12) extends ∼80 km northwestward from the Salton Buttes. The southern anomaly, which lies mostly south of the United States-Mexico border, with the lowest velocities beneath the Laguna Salada Basin, is more diﬀuse than the northern anomaly. The southern anomaly seems to extend southwestward beneath the PRB, beyond the Salton Trough (Figure 2.1b). We observe a low-velocity anomaly in the longer periods of our group-velocity model (Figure 2.4d) similar in shape to the low-velocity anomalies in our VS model (Figure 2.12). CHAPTER 2. BARAK ET AL., G-CUBED 2015 42

Interpretation and Discussion

The sedimentary basin in the Imperial Valley of the ST has been imaged as 5–6 km deep with sediments grading into metamorphic rocks (e.g., Fuis et al., 1984; Han et al., 2013), with which our model agrees. The mid-crustal high S-velocities (∼3.6–3.8 km/s) in our model beneath the ST are in good agreement with previous upper-crustal tomographic models of Lin et al. (2007) and Allam and Ben-Zion (2012), that show high VP (6.7–7.2 km/s) and high VP /VS (1.75–1.85) between 12 and 16 km depth.

The VP and VS combination implies gabbroic composition for dry rocks at this depth (Christensen, 1996). The shallowest appearance of velocities that equal or exceed 3.7 km/s at ∼12 km depth SE of the BSZ (Figures 2.11a and 2.7b) is congruent with a subbasement high observed on SSIP line 1 (Han et al., 2013). The very shallow high-velocity body we observe beneath the Salton Buttes at ∼7 km depth is also apparent in the active-source results, interpreted as a basement high (e.g., Fuis et al., 1984; Han et al., 2013). The S-velocities (∼3.6–3.7 km/s) of this body, which fit rocks of intermediate or mafic composition at this depth (Christensen, 1996) and a magnetic high at this location (Athens et al., 2011), leads us to believe that this fast anomaly most likely represents a shallow intrusion that provides heat to the overlying Salton Sea geothermal field.

The elevated velocities (VS ≥3.7 km/s) in our model appear to be conﬁned beneath the ST and are spatially separated from equally high velocities beneath the PRB to the west (Figure 2.11b). Crustal velocities do not exceed 3.7 km/s east of the Sand Hills fault system. Our result disagrees with Parsons and McCarthy (1996,

’s) interpretation of VP ∼ 6.9 km/s gabbroic lower crust extending 50 km northeast of the ST beneath the Chocolate Mountains, though their high velocities were not required by travel-times, and only constrained by amplitude modelling. Our results also do not agree with CVM-S4.26 (supporting information Figures S57) although the resolution of that model and our model decreases rapidly at the SE corner of our area of interest (supporting information S37–S41) (Lee et al., 2014).

The lower-crustal low S-velocities (VS ≤3.5 km/s) we observe beneath the ST seem surprising, because Fuis et al. (1984) interpreted VP of up to 7.2 km/s in the lower crust of the ST as gabbroic. Parsons and McCarthy (1996) used a subsequent reﬂection proﬁle to reduce the estimate of VP in the deep crust beneath the ST to CHAPTER 2. BARAK ET AL., G-CUBED 2015 43

6.9 km/s, and preliminary results from the most recent SSIP wide-angle experiment reduce the estimate again to ≤6.8 km/s (Han et al., 2013). Nicolas (1985) speculated that serpentinite was widespread beneath the ST, but this is precluded both by our S-velocities that are too high (Christensen, 1996) and by high geothermal gradients that predict temperatures too high for serpentinite to be stable (section 2.6.3). High

VP and relatively low VS also argue against the existence of older continental crust in the lower crust beneath the ST (e.g., Christensen, 1996). Conversely, our analysis below (see section 2.6.3) shows that 1% melt is enough to cause the observed velocity reduction, and up to 4.5% melt is possible, depending on the pore geometry. In addition, basaltic xenoliths from the Salton Buttes (Schmitt and Vazquez, 2006) and thermal modelling (Karakas et al., 2014) suggest melt is present today beneath the ST. The upper-mantle low-velocity we observe beneath the ST is not surprising, and has been inferred in many analyses (e.g., Parsons and McCarthy, 1996; Pollitz and Snoke, 2010; Schmandt and Humphreys, 2010; Yang and Forsyth, 2006; Yang and Ritzwoller, 2008). Although the geometry and magnitude of our upper-mantle low- velocity zone south of the Unites States-Mexico border is only poorly constrained since there we have relatively few ray paths (Figure 2.1a and supporting information Figures S23–S41), north of the border our resolution is good (supporting information Figures S23–S41), and the low-velocity zone is clearly indicated by our longest-period group-velocities (Figure 2.4d). A shallow, 40 km deep, negative converter, congruent with the location and depth of our LVZ, has been observed on Sp receiver-functions, and interpreted as thinned or new lithosphere (Lekic et al., 2011). The upper-mantle velocities we observe (in places VS <4 km/s) are also similar to low VS anomalies observed in the upper mantle along the Gulf of California (GOC) (Luccio et al., 2014;

Wang et al., 2009). In both the ST and the GOC, the top of the low-VS zone is at ∼40-km depth. Wang et al. (2009) suggested a ∼250 km spacing between low-

VS anomalies along the length of the GOC, in agreement with numerical models of buoyant mantle instabilities triggered by melting. Our low-velocity anomaly beneath the ST is indeed ∼250 km NW of their northernmost GOC anomaly, but the existence of the intervening Cerro Prieto anomaly suggests that no such simple pattern exists. As in the northernmost GOC, our upper-mantle low-velocity anomalies are oﬀset or CHAPTER 2. BARAK ET AL., G-CUBED 2015 44 extend to the west from the inferred active spreading centers (e.g., Elders et al., 1972) (Figure 2.12). Westward migration of the spreading centers has been inferred in the ST region (Parsons and McCarthy, 1996) and in the northern GOC (Arag´on-Arreola and Martin-Barajas, 2007). Finally, we note that the lowest upper-mantle velocities are located beneath the regions of lowest elevation in the ST (Salton Sea Basin and Laguna Salada Basin). This pattern also mimics the depth-to-basement shown in Fuis and Kohler (1984) and (Kohler and Fuis, 1986). This may suggest a connection between deep magmatic activity and rift-related subsidence in the ST. CHAPTER 2. BARAK ET AL., G-CUBED 2015 45

Figure 2.13: Plot of laboratory and ﬁeld measurements of VP versus VS to estimate compo- 7.4 600 MPa − Lower Crust sition and volume of partial melt beneath the ° ° ° (a) 20 /25 C 600 C 1.9 1.85 Christensen 1996 Salton Trough (a) in the crust and (b) in the up- GAB Wang Q. et al. 2013 per mantle. Field measurements: gray-shaded 7.2 rectangles show the observed range of veloci-

1.8 ties constrained by our model (VS) and active- ANO AGR MGR source data (VP ) Fuis et al. (1984); Han et al. AMP (2013); Parsons and McCarthy (1996). Red 7.0 Fuis ’84 MGR lines and arrows show the upper limit or range LVZ HVGABZ of VP for lower crust and upper mantle from

Vp [km/s] GAB 1.75 these three separate seismic experiments. In P&M ’96 MGR 6.8 the crust, LVZ: lower-crustal low-velocity zone; DIA HVZ: midcrustal high-velocity zone. In the GAB 5 Dike Dmntd. 1.7 mantle, Lid: normal velocity upper-mantle lid; 2.2 DIO

Han ’13 LVZ: upper-mantle low velocity zone. Labora- 6.6 Textural Eqlbrm. 6 1.00 5 6 1. 2.2 1. 1.0 FGR tory measurements: Squares Christensen (1996) and diamonds Wang et al. (2013a) are colored 3.4 3.6 Vs [km/s] 3.8 4.0 yellow for lower temperature measurements (20 % melt or 25◦C at 600 MPa in Figure 2.13a, 600◦C 024 at 1000 MPa in Figure 2.13b) and orange for 8.2 higher temperature (600◦C at 600 MPa in (a) (b) 1000 MPa − Upper Mantle PER and 1000◦C at 1000 MPa in (b)). Rock types: 600°C 1300°C DUN Christensen 1996 AGR: Anorthositic granulite; AMP: Amphibo- Wang Q. et al. 2013 lite; ANO: Anorthosite; DIA: Diabase; DIO: 8.0 Diorite; DUN: Dunite; FGR: Felsic granulite; 1.75 1.8 GAB: Gabbro; MGR: Maﬁc granulite; PER: 1.9 1.85 PER 1.7 Peridotite; PYX: Pyroxenite. Error bars show LVZ Lid the mean error for the data from Christensen et 7.8 al., 1996Christensen (1996) and Wang, Q. et al., Fuis et al. ’84 2013Wang et al. (2013a). In both Figures 2.13a DUN

Vp [km/s] and 2.13b, we show the eﬀect of melt on seis- Parsons & McCarthy ’96 PYX mic wavespeed for a single rock type (gabbro in 7.6 the lower crust, peridotite in the upper mantle). Lines of colored squares originating at these rock Han et al. ’13 5 Dike Dominated types (GAB and PER) show volume of par- 2.2 Textural Equilibrium tial melt along curves of constant dlnVS/dlnVP 7.4 0 4.0 1.6 1. 4.2 4.4 4.6 (tilted numbers, 1.0, 1.6, 2.25) calculated for Vs [km/s] these rock types. A line of colored squares that passes through both dark gray (HVZ, Lid) and light gray (LVZ) shaded-rectangles corresponds to a single rock composition that could represent both the high- and low-velocity lower crust, or both the lid and low-velocity upper mantle, with diﬀerent melt fractions controlling the change in velocity. Thin-dotted lines are lines of constant VP /VS. CHAPTER 2. BARAK ET AL., G-CUBED 2015 46

Estimation of Partial Melt

Magnetotellurics (MT) is often used to identify partial melt beneath the surface. However, the relatively old MT data (Martinez et al., 1989; Park et al., 1992) collected above the ST do not successfully image beneath the highly conductive zones present within the ST. Therefore, we evaluate the presence of partial melt beneath the ST using the ratio of the fractional changes in VS and VP ,“dlnVS/dlnVP ” (Takei, 2002).

We choose plausible rock compositions without melt, and calculate dlnVS/dlnVP for the range of observed P- and S- velocities beneath the ST (Figure 2.13, Appendix A).

Takei (2002) suggests dlnVS/dlnVP values of 1–1.5 are appropriate for rock+melt in textural equilibrium, and 1.6–2.25 for a region dominated by dikes and veins. Figure 2.13 shows predicted melt volume (small colored dots) for example rock types (gabbro in the crust and peridotite in the upper mantle) along lines of constant dlnVS/dlnVP values of 2.25, 1.6, and 1. To estimate the degree of partial melt in the lower crust beneath the ST, we assume a gabbroic composition, and use laboratory data from Christensen (1996) and Wang et al. (2013a) (Figure 2.13a). The laboratory data show a wide range of P- and S- velocities due to differing mineral composition: Christensen (1996)’s gabbro is gabbro- norite-troctolite whereas Wang et al. (2013a) averaged data from several sources. Our mid and lower crustal high S-velocities are quite consistent with Christensen (1996) sub solidus gabbroic composition (Figure 2.13a). In contrast, no relevant composition plots on our lower-crustal low-velocity zone (LVZ) and the addition of melt is required. Assuming our lower-crustal LVZ represents rock of the same gabbroic composition, we require <1% for a region dominated with dikes and veins and up to 4.5% melt for a system in textural equilibrium. These values are consistent with the estimate of Karakas et al. (2014) of lenses <1–km thick with 20–40% melt, if taken as an average over the whole lower crust. For the upper mantle, we assume a composition of peridotite using the laboratory results of (Wang et al., 2013a) (Figure 2.13b). Our calculation suggests <2% partial melt in the LVZ and <1% melt in the mantle lid if these regions are dominated by dikes and veins, or up to 6% melt in the LVZ and <2% melt in the lid if the upper mantle is in textural equilibrium. Humphreys and Hager (1990) estimated 1–4% melt in the upper mantle beneath the ST, based on a ∼4% reduction in VP compared to the CHAPTER 2. BARAK ET AL., G-CUBED 2015 47 average for southern California. Our assumption of uniformly gabbroic middle and lower crust, and peridotite upper-mantle, is more easily satisfied with the presence of melt in dikes that may lead to significant but localized lithospheric anisotropy (Kinsella et al., 2012).

2.7 Summary

We have presented here an improved shear-wave velocity model of southern California that extends to a depth of 50 km, using ambient-noise surface-wave tomography. Our model images new structures, mostly in the lower crust and upper mantle beneath the San Andreas Fault, Peninsular Ranges, southern Sierra Nevada, and the Salton Trough, that are mostly absent in CVM-H11.9 (Plesch et al., 2011). Our results attest to the power of ANT in illuminating lower-crustal and upper-mantle structure in southern California, where traditional event-based tomography is limited to the lateral and vertical distribution of seismicity. Across the SAF, southeast of the San Gorgonio Pass, we observe vertical to steeply- dipping lateral velocity contrasts that reach to and below the Moho. Along the San Bernardino Mountains, we observe a lower-crustal high-velocity wedge of Peninsular Ranges material below the base of seismicity, which underthrusts the SAF. We spec- ulate that the relatively shallow NE dip of the SAF (Fuis et al., 2012) in this section is not only sub-parallel with, but influenced by, this broad lower-crustal wedge that is most likely part of the arc’s mafic root, or possibly a remnant of the Farallon oceanic slab. Beneath the PRB, we observe higher crustal velocities beneath the western PRB than the eastern PRB, confirming and refining earlier potential-field models that showed that the surface west-east compositional zoning of the PRB marks a significant crustal-scale boundary. The PRB high-velocity lower crust wedges into the upper mantle beneath the SAF along the San Bernardino Mountains, and might be related to previously observed lithospheric downwelling beneath the Transverse Ranges. The transition we observe from VS ≥3.8 km/s in the lower crust of the wPRB to VS <3.7 km/s beneath the wPRB is replicated beneath the southern Sierra Nevada. In both the wSNB and wPRB it remains uncertain whether high-velocity uppermost mantle CHAPTER 2. BARAK ET AL., G-CUBED 2015 48 represents a root to these early Mesozoic arcs, or underthrust Farallon slab. Neither the ePRB nor the eSNB have high-velocity lower crust at the present day. Our velocity models cannot show whether the eSNB and ePRB ever had arclogitic roots. Finally, beneath the ST, our model details the distribution of crustal and upper- mantle magmatism in this active rift. We image a shallow (∼7 km depth) high- velocity body that we interpret as a cooling gabbroic intrusion. We observe zones of low shear-velocity in the lower crust and upper mantle, which most likely represent melt in dikes, <1% averaged over a 5–10km thick lower-crust, and <2% averaged over 10–20km thickness of upper mantle. The low VS and high VP preclude the existence of older continental crust, and support the notion that the crust beneath the Imperial Valley is fully ruptured. The upper-mantle low-velocity zones (putative zones of magma upwelling) which extend west of the active spreading centers may indicate a continuing westward migration of the driving causes of magmatism and extension in southernmost California.

2.8 Appendix: Calculation of Pore Geometry and Partial Melt

The ratio of the fractional changes in VS and VP ,“dlnVS/dlnVP ” is indicative of pore geometry, and can also be used to estimate volume of partial melt (Takei, 2002). dlnVS/dlnVP values of 1–1.5 indicate textural equilibrium of rock and melt, while values of 1.6–2.5 correspond to a region dominated with dikes and veins.

O (1 − VS/VS ) dlnVS/dlnVP = O (2.1) (1 − VP /VP ) O O where VP and VS are the seismic velocities of rock with melt and VP and VS are the velocities of the rock without melt. Rearranging equation 2.1, one can calculate O O possible pairs of VP and VS, assuming values for dlnVS/dlnVP , VP and VS :

O ! ! VP O 1 VP = O · VS + VP · 1 − (2.2) VS · (dlnVS/dlnVP ) dlnVS/dlnVP

For each pair of VP and VS, we can then estimate fraction of melt using ﬁgure

5 in Takei (2002). In Takei’s figure 5, for each value of dlnVS/dlnVP , we find the CHAPTER 2. BARAK ET AL., G-CUBED 2015 49 corresponding aspect ratio, α, which is centered between the curves representing liquid compressibility (β) values of 5 and 10, which corresponds to a rock+melt system at 0- 50 km depth (Takei, 2002). We calculate the melt fraction, using the curve in Takei’S figure 5 (top).

2.9 Acknowledgments

This study was funded by NSF-Geophysics-0911743. The seismic instruments for bb- SSIP were provided by the Incorporated Research Institutions for Seismology (IRIS) through the PASSCAL Instrument Center at New Mexico Tech. Data collected are available through the IRIS Data Management Center, that also provided us access to waveforms, related metadata, and/or derived products used in this study. The facilities of the IRIS Consortium are supported by the National Science Foundation under Cooperative Agreement EAR-1261681 and the DOE National Nuclear Security Administration. We thank CalTech for providing instrumentation and support for our stations in Mexico (M01 and M02). CICESE (Ensenada) helped with site installation in Mexico and provided data from station MBIG. The UC-Desert Research & Extension Center, Holtville, hosted our instrument center. Numerous students from Stanford, SF State, IVC, UCLA and SDSU assisted with bbSSIP station installation and maintenance. Many private landowners as well as BLM, Cal-FIRE, CDCR, State Lands Commission, USN, SD Water, and Anza-Borrego State Park allowed access to their land. We thank M. Moschetti for sharing his group-velocity model, G. Fuis for sharing his SAF model and advice, V. Langenheim for sharing data and ideas, M. Grove for geological insights, and two anonymous reviewers for their valuable sugges- tions. All ﬁgures were made using the public domain Generic Mapping Tools (GMT) software (Wessel et al., 2013). Chapter 3

Rapid variation in upper-mantle rheology across the San Andreas fault system and Salton Trough, southernmost California

An edited version of this paper was published in Geology by GSA. Copyright (2016) Geological Society of America: Barak, S. & Klemperer, S., L., Rapid variation in upper-mantle rheology across the San Andreas fault system and Salton Trough, southernmost California, USA, Ge- ology Jul 2016, 44 (7) 575-578; DOI: 10.1130/G37847.1. To view the published open abstract, go to http://dx.doi.org and enter the DOI.

Supplementary Material for this paper is available at the Stanford Digital Repos- itory: https://purl.stanford.edu/df776kg9323

50 CHAPTER 3. BARAK & KLEMPERER, GEOLOGY 2016 51

3.1 Abstract

We present new shear-wave splitting data showing systematic lateral variations in upper-mantle anisotropy across the plate boundary in southernmost California (USA). Beneath the Peninsular Ranges batholith (PRB), fast polarization directions parallel the direction of former Farallon subduction, suggestive of a slab remnant. Near the eastern edge of the PRB, across the Elsinore fault, fast polarization directions change rapidly to align with the direction of San Andreas shear. We infer that the Elsinore fault penetrates the entire lithosphere and may represent a future localization of the plate boundary that is migrating west from the San Andreas fault. Beneath the Salton Trough and the Chocolate Mountains region, large splitting times, despite a very thin lithosphere, imply vertical melt pockets in the uppermost mantle aligned in the shear direction. Largest splitting times, ∼1.2 s, are seen closest to the Sand Hills fault that projects southeast from the San Andreas fault. Further east, in the southern Basin and Range province, fast directions align with North America absolute plate motion.

3.2 Introduction

Analysis of teleseismic shear-wave splitting is a standard tool to study upper-mantle anisotropy created by strain-induced lattice-preferred orientation of minerals or by preferentially oriented melt-ﬁlled inclusions, hence also changes in rheology (e.g., Sav- age, 1999, and references therein). A shear wave passing through an anisotropic medium splits into slow and fast waves with orthogonal polarizations. Two splitting parameters (polarization, φ, and delay time, δt) provide a direct estimate of the axis and magnitude of the anisotropy for simple cases, and show systematic variations with back-azimuth to the source earthquake in more complex scenarios. Despite South- ern Californias (USA) complex tectonic history and active plate boundary oriented northwest-southeast (e.g., Dickinson, 2008, 2009; Barak et al., 2015) (Fig. 3.1A), previous studies of shear-wave splitting in Southern California (e.g., Polet and Kanamori, 2002; Kosarian et al., 2011) show a nearly uniform fast axis of anisotropy oriented approximately west-east (Fig. 3.1C) interpreted as due to inherited North America plate motion (Kosarian et al., 2011), or to pre-late Cenozoic compression (Polet and Kanamori, 2002), or to mantle ﬂow around the southern edge of the subducting Gorda CHAPTER 3. BARAK & KLEMPERER, GEOLOGY 2016 52

Barak and Klemperer - Fig. 1 - G37847

34˚N

S PAC A F ECSZ

SJ F S CMR E H BRP F F AZ 33˚N

ST CA

USA A PRB IF F Mexico L { SF USA N CPF Mexico Pacific Gulf of M7.2 N 32˚N Ocean SAFs B California 117˚W 116˚W 115˚W 114˚W −113˚ A 34˚N

NAM(HS)

33˚N ] º PAC(NAM) N 40 PAC(HS) 20 32˚N 10 mm/yr 0 C 1.0 s Diff. from SAFs [

Figure 3.1: A: Pacific (PAC)-North America (NAM) plate boundary (black: transform, orange: spreading center), showing location of B and C (black rectangle), and shear-wave splitting (SWS) fast polarizations further south in Mexico (Long, 2010) using length scale and color scale as in Figure 3.1. B: Shaded-relief topography overlain with our seismic stations (stars), other permanent stations shown in Figure 3.2 (triangles), and line of section in Figure 3.2 (white line). BRP-southern Basin and Range province; CMR-Chocolate Mountains Region; PRB-Peninsular Ranges Batholith; ST-Salton Trough. Dot-patterned orange boxes show active spreading centers (Elders et al., 1972). Strike-slip focal mechanism: 4 April 2010 El Mayor-Cucapah earthquake. Thick red lines show major faults of the San Andreas fault system (SAFs) (AF-Algodones, CPF-Cerro Prieto, EF-Elsinore, IF- Imperial, LSF-Laguna Salada, SAF-San Andreas, SHF-Sand Hills, SJF-San Jacinto). Thin red lines show southeast extension of Eastern California Shear Zone (ECSZ) (Darin and Dorsey, 2013). C: SWS fast polarizations with lengths proportional to delay time (colored bars). This study: white circles (open circles if no “good” or “fair” results were available); other SWS results: Wüstefeld et al. (2009); Liu et al. (2014). (2014; only ‘A’ quality data, weighted by their standard deviation). Color of the bars indicates difference in degrees from trend of SAFs. Black arrows are plate-motion vectors relative to the fixed-hotspot frame (Gripp and Gordon, 2002). Red line is 55 km depth contour of lithosphere-asthenosphere boundary (Lekic et al., 2011). Thin black lines are fault traces (Plesch et al., 2007). Copyright (2016) Geological Society of America. CHAPTER 3. BARAK & KLEMPERER, GEOLOGY 2016 53 slab (Zandt and Humphreys, 2008). The same general west-east pattern continues to the southern tip of Baja California (Mexico), west and east of the Gulf of Cali- fornia rift margins (Long (2010); Fig. 3.1A). Recent three-dimensional (3-D) tomographic inversion of shear-wave splitting measurements, despite the previous scarcity of data in southernmost California, has begun to suggest complicated structure, including a northwest fast axis of anisotropy in the Salton Trough (ST) (Monteiller and Chevrot, 2011) where Gulf of California ocean spreading propagates into continental crust along the San Andreas fault (Elders et al., 1972). Here we analyze shear-wave splitting measurements on our array of 48 seismometers that spanned southernmost California from A.D. 2011–2013 (Barak et al. (2015); Fig. 3.1B). We show systematic lateral variations in anisotropy between different geologic regions. Our data suggest new constraints on the underlying upper-mantle rheology, and implications for active tectonics and associated earthquake hazard.

3.3 Analysis of Shear-Wave Splitting Measurements

We use the Splitlab software (Wüstefeldet al., 2008) to determine splitting parameters (φ, δt) for SKS and SKKS phases from earthquakes with M≥6.0 at epicentral distances of 84–184◦ (see Fig. DR1 in the GSA Data Repository1). These phases, radially polarized at the core-mantle boundary, are only affected by anisotropy beneath the station, and their steep ascent through the mantle provides good lateral resolution. For each measurement we assign a quality: “good”, “fair”, or “poor” (Figs. DR2– DR5). We did not entirely discard poor measurements because a few stations have only poor measurements, yet averaged results from these stations are still consistent with adjacent stations (Fig. 3.1B). For the four permanent stations in the ST common to our study and that of Monteiller and Chevrot (2011), who used a different analysis method, the fast polarization directions match within error (Table DR2). For each station we searched for a best-fitting model with two layers of horizontal anisotropy, but also calculated a quality-weighted average φ and δt for the best-fitting one-layer model (see the Data Repository for details). CHAPTER 3. BARAK & KLEMPERER, GEOLOGY 2016 54

PRB ST CMR BRP A WSW EF SJF IF SHF ECSZ ENE φ/º 60 FAR NA(HS) 90 1.5 s 1.0 s 120 PAC(HS) 0.5 s PAC(NA) 150 SAF 1.5 1.0 δt/s 0 B 4 20 Moho 3 40 Low- km velocity zone 2 Vs [km/s] 60 1 LAB 80 V.E.=1.2 5 15 35 [SAF 23] 18 WSW EF SJF IF SHF 0 ECSZ ENE C 0 20 PAC NAM 40 former km Farallon subduction 60

80 V.E.=1.2

Barak and Klemperer - Figure 2 - G37847 and Klemperer - Figure Barak 0 100 200 300 Distance [km]

Figure 3.2: A: Our shear-wave splitting results and 10 previous measurements (Fig. 3.1B), projected onto the white line in Figure 3.1B. Thick (and thin) solid colored lines are least-squares linear ﬁts to split directions (and split times) for each group of stations (Peninsular Ranges Batholith [PRB]; Salton Trough [ST]; Chocolate Mountains Region [CMR]; southern Basin and Range province [BRP]). Error bars are quality-weighted standard deviation of polarization direction; circle diameters are proportional to split time. Thin horizontal magenta lines show Paciﬁc and North America absolute (PAC(HS), NAM(HS)) and relative (PAC(NA)) plate motions. Thick gray bars show Farallon (FAR) motion with respect to the North America plate, and strike of San Andreas fault system (EF-Elsinore, SJF-San Jacinto, IF-Imperial, SHF-Sand Hills, SAF-San Andreas fault; ECSZ- Eastern California Shear Zone). B: S-velocity model (color-scale) (Barak et al., 2015) overlain with lithosphere-asthenosphere boundary (LAB) from Sp receiver functions (white circles: Lekic et al. (2011)). Black lines are example SKS ray-paths, using the mean event azimuth, distance and depth. C: Tectonic cartoon. Numbers above fault symbols are maximum estimates of geodetic fault slip rates, mm/yr (Chuang and Johnson, 2011; Field et al., 2013). Yellow ovals are aligned melt channels. Curved arrows show upwelling mantle. Dashed lines are inferred shear zones. ⊗ and are block motions into and out of the page, size crudely scaled to velocity with respect to the EF. Copyright (2016) Geological Society of America. CHAPTER 3. BARAK & KLEMPERER, GEOLOGY 2016 55

3.4 Results

Despite modest variation in splitting parameters with back-azimuth at all our stations (Figs. DR6–DR55), our data are best modeled by a single layer with horizontal anisotropy, as also indicated by 3-D models that show a largely constant shear- wave polarization with depth in the ST region (Monteiller and Chevrot, 2011). Our one-layer model assumption is supported by the lateral consistency of our station- averaged splitting results and their correlation with other seismic observations (Lekic et al. (2011); Barak et al. (2015); Figs. 3.2, DR2). We observe systematic variations in station-averaged fast directions along our 250 km, WSW-ENE-trending array (Fig. 3.1). We divide our stations into four geographical groups, correlated with different geologic regions that are distinguished by different mean values of and t, hereafter φm and δtm (Fig. 3.2A; Fig. DR2; Table DR1). Across the Peninsular ◦ Ranges batholith (PRB) that is now part of the Pacific plate, φm ∼079 , orthogonal to the strike of the batholith, and δtm ∼0.8 s. Close to the Elsinore fault (EF), ◦ ◦ φm changes rapidly (over <10 km at the surface) from 079 outside the ST to 124 within the ST. δtm is ∼1.0 s, but δt gradually increases to 1.3 s across the ST (Figs. 3.1 and 3.2A; Fig. DR2). The 124◦ fast polarization is almost parallel to the EF (trend 130◦) and the Sand Hills fault (SHF) (trend 135◦). Northwest fast polarizations are also evident at the handful of other stations in the ST, north and south of the U.S.- Mexico border (Table DR1) (Fig. 3.1). Northeast of the SHF, across the Chocolate Mountains region (CMR) to the Colorado River and the southeastern continuation of the Eastern California Shear Zone (ECSZ), fast directions continue to rotate very slightly counterclockwise, and δt gradually decreases to 0.8 s. Across the southeast projection of the ECSZ, a second large change in fast direction occurs over <20 km at the surface, and maybe <10 km, albeit less well-constrained than the change near the EF because of our dispersed station geometry. In this eastern Basin and Range ◦ ◦ province domain, δtm ∼1.0s and φm ∼076 , sub-parallel to the 074 absolute plate motion of North America in a hot-spot reference frame [NAM(HS)] (Figs. 3.1C, 3.2A and DR2). CHAPTER 3. BARAK & KLEMPERER, GEOLOGY 2016 56

3.5 Discussion

Shear-wave splitting is commonly ascribed either to vertically coherent deformation in which anisotropy from the last significant tectonic episode recorded in the crust is also preserved in the mantle lithosphere, or to asthenospheric flow for which the associated anisotropy is governed by absolute plate motions (e.g., Savage, 1999). In our study area, we expect the common A-type (or D- or E-type) olivine fabrics that produce shear-parallel polarization of the fast directions, and not B- or C-types that require much higher water content and/or much higher stress than expected in the shallow mantle of Southern California (Karato et al., 2008). The PRB, a Mesozoic magmatic arc formed above the subducting Farallon plate (e.g., Dickinson, 2008, ◦ ◦ 2009), has φm ∼079 , oriented 34 counterclockwise from the Pacific absolute plate motion [PAC(HS) = 293◦], thereby excluding asthenospheric flow due to Pacific absolute plate motion as the cause of the anisotropy. Our observed fast directions, as well as others in the PRB north and south of the U.S.-Mexico border (Figs. 3.1A and 3.1C), are ubiquitously orthogonal to the strike of the PRB (Fig. 3.1) and are aligned with both the NAM(HS) absolute plate motion (074◦) and the Farallon-NAM convergence direction of 70–80◦ (Atwater, 1998). Although the PRB was formerly part of North America, it lay directly above the subducting Farallon plate, so would not likely have acquired anisotropy due to asthenospheric flow associated with the NAM(HS) absolute plate motion. The 70–80-km-thick lithosphere of the PRB (Lekic et al., 2011) (Fig. 3.22B) is sufficient to develop the observed 0.8 s delay time, under the assumption of an ∼10 km mid-lower crustal layer with average anisotropy ∼10% (Porter et al., 2011) and up to 5% anisotropy in the mantle lithosphere, as commonly found in mantle xenoliths (Savage, 1999, e.g.,). Hence frozen-in lithospheric fabric aligned in the direction of subduction is the likely cause of the dominant west-east fast direction observed along the entire 1200 km length of the PRB (Fig. 3.1A). It is controversial whether the high-velocity upper-mantle and lower-crust of the PRB represents the magmatic arc or underthrust Farallon slab (Barak et al., 2015). The high anisotropy and relatively low upper-mantle shear velocities (Barak et al., 2015) are both more consistent with peridotite than with eclogite (Worthington et al., 2013), so we favor the interpretation of a slab remnant (cf. Wang et al., 2013b). The 6 m.y. CHAPTER 3. BARAK & KLEMPERER, GEOLOGY 2016 57 elapsed since the transform plate boundary migrated inboard of Baja California to the modern San Andreas fault system (SAFs), transferring the PRB to the Pacific plate (Stock and Hodges, 1989), must be insufficient - as modeled by Savage et al. (2004) - to have rotated an older west-east anisotropic fabric to be parallel to the SAFs or to the Pacific absolute plate motion (Fig. 3.1B). At the eastern edge of the PRB, fast directions rapidly rotate to a northwest-southeast direction over a distance of <10 km. Our 120-km-wide region of unexpected northwest-oriented fast S-wave polarizations in the ST correlates spatially with a region of upwelling magma (Barak et al., 2015) beneath a lithosphere as thin as 40 km (Lekic et al., 2011) (Fig. 3.2B). The thin lithosphere is bounded by rapid lateral increases in lithospheric thickness coincident with the boundaries implied by our splitting measurements, close to the EF and the southern extension of the ECSZ (Fig. 3.1B). Inward mantle flow (SAF-perpendicular) due to lithospheric thinning beneath the ST would produce a fast direction orthogonal to the north-northeast-elongate ST and Gulf of California, which we do not observe. Instead, the northwest fast axis-parallel to the SAFs - suggests that mantle flow beneath the rift is dominated by shearing motion in this oblique-transform plate boundary, as also expressed in the crust by the very long transform faults connecting the very short spreading centers (Elders et al., 1972) (Fig. 3.1A). The rapid lateral variation in anisotropy confines the cause of anisotropy to shallow levels, certainly above 90 km depth given our seismic waves of ∼8 s period (e.g., Hammond et al., 2010). Nonetheless, despite the shallow source, we observe ∼1.0 s of splitting delay time (Fig. 3.22A; Fig. DR2). Crustal anisotropy typically contributes only 0.1–0.3 s (e.g., Savage, 1999). Even the 1% melt organized in lower-crustal melt channels that is permitted by our velocity tomography (Barak et al., 2015) only contributes 0.25 s splitting for an aspect ratio of 0.01 (Ayele et al., 2004). The mantle lithosphere beneath the ST is so thin, just 10–15 km thick beneath the region of largest splitting times (Lekic et al., 2011; Barak et al., 2015) (Figs. 3.2B and 3.1C), that shearing would only produce <0.2 s splitting given typical anisotropy of sheared mantle xenoliths <5% (e.g., Savage, 1999). Therefore, to explain the large split times, we argue that vertical melt-filled channels must be present in the mantle (Ayele et al., 2004; Hammond et al., 2010; Holtzman and Kendall, 2010) aligned in the transform shear direction (northwest-southeast) and orthogonal to the direction of least-principal stress. CHAPTER 3. BARAK & KLEMPERER, GEOLOGY 2016 58

Oblique rift systems can exhibit a difference in dike orientation between the crust and the mantle (Abelson and Agnon, 1997). Injection of melt into the crust over short elastic time scales results in dikes oriented orthogonal to the shear plane to minimize shear on their walls. In contrast, upper-mantle melt conduits form continuously in a ductile flow regime that reduces effective viscosity (Holtzman and Kendall, 2010), re- laxes lithospheric stresses, and results in shear-parallel dikes regardless of the oblique divergence of the overlying plates (Abelson and Agnon, 1997). Assuming the onset of melt occurs at depths of ∼70 km (Wang et al., 2013b; Lekic et al., 2011), a 30- km-thick layer with 1% melt organized in dikes, as inferred from our tomography (Barak et al., 2015), contributes 1 s split time for a 0.01 aspect ratio (Ayele et al., 2004). The rapid change in fast direction across the EF can be modeled as due to a laterally varying viscosity structure (Savage et al., 2004; Bonnin et al., 2012) in which the greatest strain is concentrated within a low-viscosity region to the east (Pollitz et al., 2012). Because our SKS raypaths are mostly from the west and southwest, our stations are mapping anisotropic fabric development up to 25 km southwest of the EF at 50 km depth, within the mantle lithosphere (Fig. 3.2B). There is no evidence that the EF dips southwest through the crust (Plesch et al., 2007; Barak et al., 2015), but our velocity model (Fig. 3.2B) and common conversion-point image (Fig. DR2) show that the upper-mantle low-velocity zone extends west beneath the EF, suggesting gradual heating of the cold PRB/Pacific plate by hot ST mantle. Because western strands of the SAFs are younger (EF and SJF: 1.1–1.3Ma, Dorsey et al. (2012); SAF > 5.5 Ma, Stock and Hodges (1989)), we suggest the plate boundary is migrating west to the EF (Fig. 3.2C). Westward migration has been previously inferred in the ST (Barak et al., 2015) and the northern Gulf of California (Aragón-Arreolaand Martin-Barajas, 2007). Mean splitting parameters barely change from the ST to the ◦ ◦ CMR (ST: φm ∼124 , δtm ∼1.0 s; CMR: φm ∼117 , δtm ∼0.9 s). However, these averages mask a significant change in split times, from δt ∼0.9 s at the EF to δt ∼1.3 s at the SHF, then a decrease back to δt ∼0.8 s at the ECSZ (Fig. 3.2A; Fig. DR2). Because the lithosphere is almost as thin beneath the CMR as beneath the ST (Fig. 3.2B), there must still be a large contribution from the asthenosphere northeast of the SHF to account for the 0.9 s delay time. Hence, vertical melt inclusions likely also exist beneath the CMR, largely rotated into alignment with the plate-boundary CHAPTER 3. BARAK & KLEMPERER, GEOLOGY 2016 59

SAFs shearing direction. The maximum split times are beneath the SHF, the projection of the SAF that was the longest-lived locus of plate-boundary transform motion within the SAFs. A similar magnitude of decreasing delay times away from the SAF has been modeled in northern California (Bonnin et al., 2012). Within the southern ◦ Basin and Range province, φm (∼076 ) is aligned with the local NAM(HS) absolute plate motion (074◦), suggesting shear at the lithosphere-asthenosphere boundary likely complemented by extensional vertically coherent deformation.

3.6 Conclusions

The distributed plate boundary in southernmost California coincides with an unusually weak uppermost mantle of melt-lubricated asthenosphere. Thermally controlled upper-mantle viscosity helps to localize deformation and strongly aﬀects brittle deformation at the surface (Molnar, 1992). A large discrepancy between slip rates calculated using one-dimensional viscosity models and those measured by GPS stations and geologic data in and around the ST is unparalleled elsewhere in California (Chuang and Johnson, 2011; Field et al., 2013). The rapid lateral changes in viscosity we infer across the EF and the ECSZ, previously only conjectured from satellite geodesy (Pollitz et al., 2012), should inform future inversions of slip rate, estimates of earthquake hazard, and understanding of geologic evolution.

3.7 Acknowledgments

We thank C. Castillo, S. Chevrot, V. Langenheim, K. Liu, W. Thatcher and an anonymous reviewer for valuable comments. Our seismic data are available from http://ds.iris.edu/ds/nodes/dmc/. This work was funded by National Science Foun- dation Geophysics program grant 0911743. Chapter 4

Imaging upper-mantle upwelling beneath the Salton Trough using joint inversion of receiver functions and surface wave dispersion curves

Please see Word document. Supplementary Material for this chapter is available at the Stanford Digital Repos- itory: https://purl.stanford.edu/df776kg9323

60 CHAPTER 4. JOINT INVERSION OF RF AND SWD 61

Supplementary Material for this paper is available at the Stanford Digital Repos- itory: https://purl.stanford.edu/df776kg9323 CHAPTER 4. JOINT INVERSION OF RF AND SWD 62

4.1 Abstract

4.2 Introduction

4.3 Data and Method

4.4 Results

4.5 Discussion

4.6 Conclusions

4.7 Acknowledgments Appendix A

Data Acquisition

In January 2011, I deployed 42 broadband seismometers across the Salton Trough, provided by the PASSCAL pool and Caltech, as part of the Salton Seismic Imaging Project (SSIP). 36 of the 42 stations were deployed at 5-10 km intervals, along a line perpendicular to the rift axis. The other six stations, including two in Mexico, were deployed off-line to provide areal coverage. This array scheme was chosen to allow for multiple crossing ray-paths in the crust and upper mantle, important both for Common-conversion-point migration of receiver functions and ambient noise tomography. Stations were nominally deployed for 18 months in order to provide a good ensemble of large earthquakes with good azimuthal distribution, necessary for receiver function analysis. Actual recording time varied between 6 to 24 months. Our main line coincides with older and recent refraction profiles Fuis et al. (1984); Parsons and McCarthy (1996) and the more recent SSIP controlled-source profile, acquired in March 2011, both across and along the STRose et al. (2013); Persaud et al. (2016). The controlled-source part of the SSIP was lead by John Hole, Joanne Stock and Gary Fuis. The acquisition process, including scouting, permitting, deployment and service, has proved to be a big challenge, but a great learning experience for me. Site locations along the profile were chosen to allow vehicle access while maximizing security and minimizing potential noise sources. Given these limitations, our various site locations included privately owned lands, corporate-owned lands, prisons, fire stations, and public land. Permitting and field logistics involved working with both private and

63 APPENDIX A. DATA ACQUISITION 64 corporate land owners, and several agencies such as BLM, CA SLC, CDCR, US Navy, etc. Due to permitting issues, three of the sites (A15, A30, A33) were installed late in March 2011 and another site (A21) was installed only in June 2011. Sensors intended for these sites were installed along a smaller cross-line profile (Figure xxx). We also had to move two of the sites, A26 (to A26A) and A07 (to A07A), due to proximity to a busy highway and equipment theft, respectively. The instrument pool included 6 different types of sensors (STS-2, Guralp ESP, Gu- ralp 3T, Trillium 40, Trillium 120 and Trillium Compact). The installation involved two distinct deployment strategies: 3Ts were buried directly in the ground while the other sensors were buried in vaults, which had to be customized to accommodate the different sensor types. The installation crew included two technicians from PASSS- CAL, and 15 students from four different universities. Our lodging and instrument center was established at the UC-Davis Desert Research and Extension Center in El Centro, CA. After the installation, I returned to service the stations every quarter (9 service-runs in total). After each service-run, I composed and sent to PASSCAL detailed reports of issues we encountered. All the data has been archived in the IRIS DMC and is now publicly available through IRIS. Various problems with the equipment lead to significant data loss. The main cause of data loss was a phenomenon of “mass derailing” in the STS-2 sensors in response to automatic centering pulses. The “mass derailing” I have mitigated this problem by increasing the frequency of auto re-centering from a 6-day to a 1-day interval, and by replacing some of the sensors. Still, this phenomenon is responsible of about 50% data loss in some of our stations. Another cause of data loss are gaps in recording. These recording gaps amount to a total of ∼23 months, from which ∼82% (18 months) was due to electrical and mechanical malfunctions, which included: data-logger malfunctions (freezing, full-scale noise, disk writing problems), sensor malfunctions, and power loss due to overheating of power boxes. The remaining 18% included equipment theft, power loss due to a loose battery connection, and a solar power cable eaten by wildlife. Quality analysis of the recorded data revealed high-frequency noise in the stations located in the basin, probably cultural noise (mainly traffic). In addition, long-period noise observed in data recorded at some of the Guralp ESP stations is suggested by PASSCAL to be due to fluctuations in APPENDIX A. DATA ACQUISITION 65 barometric pressure and temperature, due to shallow burial of the vaults. In addition, towards the end of the experiment we discovered that in several Guralp ESP sensors the data and power cable was bent by pressure from the lid of the vault, which caved in under the pressure of the dirt piled above. The bending of the cable caused distortion of the recorded signal. In some stations, damage to the cable was observed visually. To resolve this issue we replaced the initial vaults with higher vaults where it was necessary. Because of the significant amount of data lost or corrupted during the allotted time period of the experiment, we received authorization to extend the experiment until March 2011, with only the ESP sensors. In the additional ∼6 months of recording the ESP sensors were deployed, in the new tall vaults, at the previous site locations in the Imperial Valley, since these sites suffered the most data loss and noise. Bibliography

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