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CALIFORNIA STATE UNIVERSITY, NORTHRIDGE RAYLEIGH WAVE TOMOGRAPHY BENEATH THE OCEANIC AND CONTINENTAL MARGIN OF THE NORTH AMERICAN AND PACIFIC PLATE BOUNDARY A thesis submitted in partial fulfillment of the requirements For the degree of Master of Science in Geology, Geophysics By Sampath Chaminda Bandara Rathnayaka Mudiyanselage December 2014 The thesis of Sampath Chaminda Bandara Rathnayaka Mudiyanselage is approved: Dr. Monica Kohler Date Dr. Gerry Simila Date Dr. Dayanthie Weeraratne, Chair Date California State University, Northridge ii TABLE OF CONTENTS SIGNATURE PAGE………………………………………………………………………ii TABLE OF CONTENTS ................................................................................................... iii List of Figures ..................................................................................................................... v List of Tables...................................................................................................................... vi ABSTRACT...................................................................................................................... vii Chapter 1 Introduction ........................................................................................................ 1 1.1 Motivation ................................................................................................................. 1 1.2. Tectonic history ........................................................................................................ 2 1.3. Previous work deep seafloor ...................................................................................11 1.4. Previous work done by Escobar at el.,2013 ........................................................... 15 Chapter 2 Seismic Data..................................................................................................... 16 2.1 OBS Deployment and Retrieval.............................................................................. 16 2.2 Land Stations .......................................................................................................... 18 2.3 Data Corrections ..................................................................................................... 18 2.4 Seismic events ......................................................................................................... 25 Chapter 3 Surface Wave Method ...................................................................................... 28 Chapter 4 Results .............................................................................................................. 31 4.1 Regional Phase velocity results .............................................................................. 31 iii 4.2 Regional Anisotropy results ................................................................................... 39 Chapter 5 Discussion .................................................................................................... 48 5.1 Deep seafloor .......................................................................................................... 48 5.2 Ocean versus continetal lithosphere........................................................................ 48 5.3 The Borderland ....................................................................................................... 52 Chapter 6 Conclusion........................................................................................................ 55 References ......................................................................................................................... 56 Appendix A : Amplitude correction plots ........................................................................ 61 Appendix B: Earthquakes ................................................................................................ 73 Appendix C: Station Correction........................................................................................ 77 Appendix D: Phase Velocities .......................................................................................... 78 Appendix E: Seismograms ............................................................................................... 79 iv List of Figures Figure 1.1 Simplified map of the Borderlands 5 Figure 1.2 Crust and Lithospheric variation across my study area 8 Figure 1.3 Seafloor magnetic isochron map 14 Figure 2.1 ALBACORE OBS deployment area 16 Figure 2.2 Behavior of amplitudes with increasing depths 20 Figure 2.3 S-Wave velocity variation with depth 22 Figure 2.4 Phase Velocity variation with period (Starting model) 23 Figure 2.5 Behavior of amplitudes after applying amplitude correction 24 Figure 2.6 Azimuthal distribution of earthquakes 26 Figure 2.7 Ray path coverage for the study area 27 Figure 3.1 Grid node configuration 30 Figure 4.1 Average phase velocity as a function of period in three regions 32 Figure 4.2 Average phase velocity as a function of period in five regions 35 Figure 4.3 Average phase velocity as a function of period in seven regions 37 Figure 4.4 Anisotropy deep seafloor 40 Figure 4.5 Anisotropy mid age seafloor 40 Figure 4.6 Anisotropy outer Borderland 42 Figure 4.7 Anisotropy inner Borderland 42 Figure 4.8 Anisotropy North Garlock region 43 Figure 4.9 Anisotropy South Garlock region 44 Figure 4.10 Anisotropy West San Andreas region 45 v List of Tables Table 2.1 P-wave, S-wave, and density variation in study region vi 21 ABSTRACT RAYLEIGH WAVE TOMOGRAPHY BENEATH THE OCEANIC AND CONTINENTAL MARGIN OF THE NORTH AMERICAN AND PACIFIC PLATE BOUNDARY BY Sampath Chaminda Bandara Rathnayaka Mudiyanselage Master of Science in Geology, Geophysics The North American and Pacific plates are separated by a unique transform plate boundary in southern California. The inception of the San Andreas fault system formed as a result of subduction of the East Pacific Rise spreading center, rifting of the Borderlands in the Miocene, and subsequent plate rotation that is ongoing today. However, the stresses surrounding this tectonic system are only partially understood due to the lack of offshore data which makes up half of this plate boundary. I used Rayleigh waves recorded by a marine seismic array of 34 ocean bottom seismometers (OBS) deployed as part of the ALBACORE (Asthenospheric and Lithospheric Broadband Architecture from the California Offshore Region Experiment) project offshore southern California on 18-32 Ma seafloor. The marine seismic array recorded data for a 12 month duration from August 2010 to 2011, and are combined with 82 land stations from the CISN network which recorded earthquake data simultaneously. I analyzed 80 teleseismic events at distances ranging from 30° to 120° for Mw ≥ 5.9, filtered at periods between 16 and 78 s. Strong gradients in water depth, sediments, and crustal thickness are present across this plate margin; therefore, I performed amplitude corrections for OBS stations vii that account for velocity variations in water, sediment layer, crustal thickness, marine fossil layers, and lithospheric thickness as a function of sea floor age. I used a surface wave inversion that considers a two plane wave method to represent the incoming wave field and performed a grid search for inversion parameters. My results indicated that phase velocities averaged over my study area offshore are 1.6% lower than previous studies for the seafloor age bin 20-52 Ma that used oceanic ray paths recorded by land stations only. Phase velocities in the Borderlands are lower than the deep seafloor, but higher than the southern California land region at all periods. Phase velocities at lithospheric depths are 1.4% higher in the oceanic mantle compared to the continental mantle indicating compositional and structural differences due to formation history in the two tectonic environments. Anisotropy in the Southern California land region is very uniform at 1.5 % in an average direction N 69˚ W. The inner Borderland demonstrates different anisotropic structure compared to the outer Borderland at all periods. The fast directions in the inner Borderland are NW-SE below 50 s and change to NE-SW above 50 s. Anisotropy in the outer Borderland and deep seafloor displays an E-W fast direction at short periods that is perpendicular to seafloor magnetic anomalies, and consistent with the remnant fossil spreading direction. A transition is observed at long periods above 50s in the outer Borderland and deep seafloor that has a NW-SE fast direction which is parallel to active Pacific plate motion today. This change in anisotropic fabric from fossil spreading 15-35 Ma to active plate motion today will also constrain the base of the lithosphere. viii Chapter 1 Introduction 1.1 Motivation Southern California is the location of a unique transform plate boundary on land that divides the North American plate and the Pacific plate. This tectonic system was formed during the Miocene period with the subduction of the Farallon plate and the East Pacific Rise (EPR) beneath California. The Farallon plate was subducting steadily beneath the North American plate from 37 to 30 Ma. The EPR spreading center approached the western continental boundary of North America because the subduction rate was faster than the spreading rate. The EPR subducted beneath southern California 30 Ma. After subduction of the EPR, rotation of land and Borderland blocks in a counter wise (CW) sense began 29 Ma ago (Atwater, 1989). The Borderland, Los Angeles region and Mojave Desert subsequently underwent crustal extension and rifting from 24-18 Ma ago (Wright, 1991), accommodated by normal faulting (Crouch and Suppe, 1993). The formation of the Borderland is thought to be accompanied by ~100º of ongoing counter wise (CW) rotation of the Transverse Ranges that we see today (Atwater, 1989; Luyendyk, 1991; Crouch and Suppe, 1993). This tectonic event also marked a change from a convergent zone to a transform boundary (Atwater, 1970), and created one of the few examples of a transform plate boundary (San Andreas Fault - SAF) on land. Much work has been done to understand the tectonic evolution and deformation on the continental side of the San Andreas Fault and the southern California region which is now a highly populated urban region including the Los Angeles metropolitan area. However, very little is known about the offshore side of this plate boundary due to limited and difficult access to the seafloor. Thus tectonic events such as plate-scale 1 deformation and onshore-offshore fault stresses are not well understood across the plate boundary. Here, I present seismic results from a marine deployment of 34 ocean bottom seismometers offshore southern California, the ALBACORE (Asthenospheric and Lithospheric Broadband Architecture from the California Offshore Region Experiment) project, to study the west and east side of the Pacific-North America plate boundary simultaneously. Recent studies (Legg et al., 2012) from new bathymetry data offshore (from the ALBACORE project) indicate that there are at least two major sets of faults offshore that mimic the shape and bend of the SAF, and suggest seismic hazards offshore are significant. My results indicate that phase velocities at lithospheric depths are 1.4% higher in the oceanic mantle compared to the continental mantle indicating compositional and structural differences due to formation history in the two tectonic environments. Plate structure offshore is consistent with anisotropic fabric formed during plate formation growth and conductive cooling expected for oceanic lithosphere that is perpendicular to N-S magnetic anomalies at the oldest ages in our study area (25-35 Ma). Local rotation of magnetic anomalies surrounded by arcuate pseudo-faults within the Patton and Arguello region (15-25 Ma seafloor) are reflected in a change of anisotropic fabric in a perpendicular direction at periods below 50 s. The remnant fossil anisotropy changes for periods above 50 s (~56 km depth peak sensitivity) in all regions except the inner Borderlands are consistent with NW Pacific motion direction today. 1.2 Tectonic history The tectonic reconstruction of the seafloor was conducted in previous work looking at magnetic lineations using hot spots as a stationary reference frame (Engebretson et al., 1984, Atwater, 1989). At about 110 Ma ago, the East Pacific Rise was located in the 2 central Pacific ocean forming the Pacific and Farallon plates to either side (Atwater 1989). The Isanagi and the Farallon plates moved away from the Pacific plate and the oldest edge of the Farallon plate began subducting beneath the western margin of North America. Subduction proceeded at a rate faster than spreading of the EPR, causing eastward migration of the spreading center (Atwater, 1989). The Pacific plate thus expanded with time while the Farallon plate shorted due to subduction. The Isanagi plate disappeared entirely after 70 Ma. In the late Cretaceous, about 80 Ma ago, collision and subduction of an oceanic plateau (Liu at el., 2008), caused the Farallon to form a micro plate, the Kula plate, in the north Pacific seafloor which began to subduct beneath northern Canada and the Aleutian island trench (Atwater, 1989). And in the middle Eocene, the Kula plate disappeared into the subduction zone. Throughout the Paleocene period, the Farallon plate continued moving to the east until it experienced its first disruption documented by the abrupt change in direction of the Surveyor, Mendocino and Pioneer fracture zones at about 55 Ma (Atwater, 1989). This change has been interpreted as the breakup of the northern portion of the Farallon (the region between the Pioneer and Murray fracture zone) which continued to migrate to the north and later became the Vancouver plate (Atwater, 1989). During the late Eocene, about 37 Ma ago, the western side of the North American plate was reorganized. The EPR spreading center reached the trench of the western margin of the North American plate near Baja and southern California and began subducting (Atwater, 1989). This event also initiated a transform San Andreas fault on land (Atwater, 1989). During the mid-Miocene period (at about 17 Ma), the Western Transverse Ranges (WTR) and bordering areas began to rotate in a clockwise direction. 3 Extensional features are observed due to dextral shear (transtension and transpression) between the Pacific and North American plates (Luyendyk, 1991). The total rotation is at least 93o indicating that the axis of the WTR was originally N-S but is oriented in EW direction today. This Pacific–North American deformation zone began contracting in width during Pliocene time, but clockwise rotation progresses today (Luyendyk, 1991). However, it is not clear whether the WTR continues offshore to the south or west across the coastline, or how far offshore rotational stresses are active today During Cretaceous and Paleogene time, more than 10,000 km of the Farallon oceanic plate was subducted beneath western North America, resulting in a continental margin arc-trench system (Crouch and Suppe, 1993). And this active arc-trench system included five major lithotectonic belts such as the Franciscan accretionary wedge, the underlying Coast Range ophiolites, the Great Valley forearc-basin sequence, the accreted arcs and mélanges of the Western Foothills belt, and the magmatic arc of the Sierra NevadaPeninsular ranges (Crouch and Suppe, 1993). Disruption of central California took place in the Cretaceous, during the subduction event including rotation of the WTR, opening of the Los Angeles basin and the Borderland region offshore southern California (Atwater, 1970, 1989). The California Borderland mainly consists of four geologic belts (Figure 1.1) namely the western Transverse Ranges block, the Patton accretionary belt, Nicholas forearc belt (outer Borderland), and the Catalina Schist belt (the inner Borderland) (Bohannon and Geist, 1998). 4 Figure 1.1: Simplified map of the main geologic structures found in the California Borderland. (Bohannon and Geist,1998). The kinematic evolution of California’s Borderland began in the early Neogene/ early Miocene after the subduction of the Farallon plate beneath the southern California (Atwater, 1989). At this instance, the relative motion between the Pacific and North American plates over the transform boundary had induced horizontally transmitted stress from one plate to another across a wide zone of deformation including the Borderland (Atwater, 1970). During the middle Miocene, the Monterrey fragment was captured by the Pacific plate and then Pacific-Monterrey spreading slowed and eventually ceased under California. Later the Arguello and Patton fragments were also captured by the Pacific plate (Nicholson et al., 1994), and rotation of these blocks began forming the acruate shape of these pseudo fracture zones. It is not clear, whether these rotational features are restricted to magnetic anomalies on the surface of the seafloor or if they permeate to deeper depths producing lithospheric scale deformation, extension, or rifting. 5 The Catalina Schist was uplifted from mid-crustal depths and exposed during a major event of extensional tectonism that had started in early-middle Miocene time along with about 100o of clockwise rotation of the WTR belt causing the displacement of the gently deformed Nicolas forearc belt to the west (Luyendyk, 1991; Bohannon and Geist, 1998). The rotation of the WTR and the translation of the Nicolas forearc belts were associated with a large amount of uplift that probably involved strong flexural deformation of the footwall (Bohannon and Geist, 1998). Extension in the Borderland continued during and after middle Miocene and as a result, most of the Borderland was further deformed into numerous ridges and basins during this stage of oblique extension (Bohannon and Geist, 1998). There are two major transpressional fault zones in the Borderland namely Ferrelo and San Clemente. Both are right lateral strike slip faults (with right –stepping en echelon pattern) trending towards northwest. This fault zones manifest by dextral offset of large scale geomorphic feature that were created during middle Miocene and later tectonic evolution of the California Borderland. New bathymetry collected during the ALBACORE cruises added new swaths in the Borderlands to map these faults. The Ferrelo and San Clemente faults may mimic the shape and bend in the San Andreas fault (Legg et al., 2014). Much work has been done to determine the crustal thickness and upper mantle structure as well as current mantle flow patterns beneath the continental side of the plate boundary in southern California. Studies which have used iterative thermo mechanical models (Gilbert et al., 2012; Saleeby et al., 2012) predict that the process of delamination of the arclogite root of the Sierra Nevada batholith is still in progress. On the other hand, crustal thickness has been estimated at 40 km beneath the San Gabriel Mountains using 6 teleseismic travel time residuals from the LARSE experiment (Kohler and Davis, 1997; Zhu and Kanamori, 2000). Results from modeling of teleseismic receiver function estimated a crustal thickness of 29 km beneath the central Transverse Ranges. A recent study, which has used the P to S receiver function method, indicate that average Moho depth below Los Angeles and the WTR is 28.5 km, 27.9 km respectively (Reeves et al., 2014). Also this study has shown that the average Moho depth below the inner Borderland and outer Borderland is 21.86 km, 28.17 km respectively. The inner Borderland (IB) was found to have a crustal thickness of 19 to 23 km (ten Brink et al., 2000) which is consistent with recent P to S conversion studies (Reeves et al., 2014). The analysis of teleseismic P waves with a 5200-station array in long beach, California (Brandon at al., 2013) has postulated a sharp northeastward increase in Moho depth from the IB to main land southern California. That Moho dips 65˚ to the northeast and flattens ~10 km southwest of the Newport-Inglewood fault zone. The crustal thickness at the continental margin beneath the Los Angeles basin decreases rapidly over a short distance from 30 km under the Transverse Ranges to about 20 km beneath the IB (Zhu and Kanamori, 2000). Lithospheric thickness was estimated recently by surface waves studies (Yang and Forsyth, 2006a) showing that lithospheric thickness beneath southern California is approximately 90 km. Studies from P wave travel time inversions indicate that the thickness of the lithosphere under the San Gabriel Mountains and San Andreas fault is 60-80 km (Kohler, 1999). A recent study, which used scattering of teleseismic shear waves, indicated important variations in the lithosphere-asthenosphere boundary (LAB) (Lekic et al., 2011). The study revealed an abrupt change in lithospheric thickness from 7 approximately 70 km beneath the Los Angeles Basin to 50 km under the inner Borderland. This study also imaged a thick lithosphere of approximately 90 km for the outer Borderland. The Moho and LAB results from the studies described are summarized in the sketch of Figure 1.2. Figure 1.2: Crust and lithospheric thicknesses schematic for my study area along a line from west to east at 33o latitude. Solid lines indicate known depths and dashed lines with question mark indicate unconfirmed depths (the location of dashed lines are based on some studies mentioned in the text) The blue color region indicates water and the brown color indicates underplating or regions which are above sea level such as sediment layers. Plate rotation was identified in the Western Transverse Ranges by using paleomagnetic observations (Luyndyk et al., 1991). Seismic studies which have used inversion of teleseismic P-wave travel-time residuals have found evidence for clockwise rotation of a high velocity anomaly with increasing depth. This begins from an EW orientation at a depth of 50 km and rotates to NNE-SSW at a depth of 190 km (Kohler et al., 2003). Seismic anisotropy studies using SKS splitting measurements have indicated 8 that the fast polarizations directions for stations east of the San Andreas (including the Mojave Desert area) are roughly E-W ( 80-95o east of north) (Liu et al., 1995). Fast polarization directions for stations west of the San Andreas (including regions south of the Great Valley) were appeared to be shifted to a NE-SW azimuth (approximately 53o east of north) over the southern end of the Great Valley. Anisotropy in the WTR and the northern Peninsular Ranges is N 82o E, N 94o E for east of north in the Mojave Desert, and N 70o E on San Clemente Island. Another study using SKS splitting measurements has shown that fast polarization directions were approximately E-W across a broad area in southern California (Polet and Kanamori, 2002) and that results were consistent with an averaged azimuthal anisotropy from surface wave studies in southern California (Yang and Forsyth, 2006a). Near the Santa Barbara bay area, the fast velocities exhibited EW direction. However, fast velocities exhibit a NW-SE direction in the region of the Channel Islands off shore (Polet and Kanamori, 2002). For Southern California, most previous studies have found that fast directions in SKS splitting measurements are dominantly ENE-WSW (Liu et al., 1995; Polet and Kanamori, 2002; Silver and Holt, 2002). Recent studies suggested that this fast direction in SKS splitting is most likely due to the strain-induced lattice-preferred orientation (LPO) of olivine (Kosarian et al., 2011). A recent combined study of SKS and surface wave splitting predicted N 112o W from Rayleigh wave fast directions for the continental lithosphere (Southern California) in a depth range of 33-100 km (Kosarian et al., 2011) which is consistent with the ENE-WSW observation from SKS splitting (Liu et al., 1995; Polet and Kanamori, 2002; Silver and Holt, 2002). According to this finding, the direction of anisotropy from surface waves is approximately parallel to the San Andreas Fault, a result that contrasts with the N 80o E 9 fast directions from the same study. The study suggested that at least two layers of anisotropy were required to explain the contrast between the results, the first layer is in the depth range of 33-100 km (mantle lithosphere) and the second layer is in a depth range of 300-400 km. The contrasting pattern may suggest, therefore, that anisotropy for southern California twists in a counter clockwise direction with increasing depth (Kosarian et al., 2011). The upper mantle flow field beneath southern and northern California has been suggested to be simply aligned with plate motion in some studies but found to be complex and non-standard in other studies. Silver and Holt (2002) which used surface motion, mantle deformation data from Global Positioning System (GPS) and seismic anisotropy respectively, initiate that mantle flow is only weakly coupled to the motion of the surface plate which produces a small drag force. Seismic tomography, anisotropy, studies was compared to kinematic plate reconstruction, showing that a toroidal mantle flow is present beneath a wide region of western North America (Zandt and Humphreys, 2008). The toroidal mantle flow field spans ~1000 km in diameter including Nevada, Utah, Arizona, southern Oregon, and California indicated it may be driven by plume flow around the plate edge or remnants of Farallon plate subduction. This toroidal pattern predicts flow in the NW-SE direction offshore southern California. Whether plate velocity is dictated by absolute plate motion of the Pacific plate (55˚ NW), the North America plate (63˚SW) or rotation, studies of shear traction suggests that the lower mantle will respond actively to mantle drag (e.g. Liu and Bird, 2002). Other studies have indicated that the anisotropy pattern which was observed in southern California (Land) is a direct result of the alignment of the a-axis of olivine crystals (E-W fast directions), 10 nearly perpendicular to the World Stress Map vectors (N-S direction) of lithospheric shortening (Polet and Kanamori, 2002). However, Kosarian et al., (2011) suggests that SKS splitting is due to drag in the asthenosphere by the plate motion of the overriding plates. The NW-SE Pacific plate motion direction (N 55o W) is roughly consistent with the toroidal flow model offshore. 1.3. Previous work in deep seafloor Based on Hess (1962) study, the oceanic plate region in our study area, is part of the Pacific basin, and was formed due to normal growth and slow cooling during propagation away from the East Pacific Rise (EPR) spreading center. The oceanic lithosphere was formed by the upwelling and partial melting of material that rises from the asthenosphere at the oceanic spreading center. This material cools and grows in thickness as it spreads away from the ridge. There are two basic models that predict the variation in depth and heat flow with increasing age as oceanic lithosphere cools. The simple cooling model, states that the lithosphere behaves as the cold upper boundary layer of a cooling half space which predicts a linear relationship between depth and age1/2, or heat flow and age1/2 of the ocean floor (Turcotte and Schubert, 2002). The Plate model, states that the lithosphere is treated as a cooling plate with an isothermal lower boundary. Here it cools conductively until it reaches a predetermined limit. According to this model, the plate thickness has an asymptotic thermal thickness of 125 km for the Pacific oceanic lithosphere (Parsons and Slater, 1977; Stein and Stein, 1992). According to the conductive cooling models, the seafloor subsides due to growth and increased density of the lithosphere as it cools. However seafloor at 80 Ma or older indicate that the data for bathymetry, gravity and heat flow show a higher values from this prediction. Small scale 11 convections cause this departure from prediction and my study area fit into region with seafloor age less than 80 Ma in this conductive cooling model. There are several studies which have determined the variation in phase velocities with age in the seafloor. The MELT experiment, for example, indicates that the average phase velocity for the youngest sea floor, less than 4-5 Ma, is two times larger than the average phase velocity in the next age zone ( age 4 – 20 Ma) and this could be explained by conductive cooling of the mantle (Harmon et al., 2009). Rayleigh wave studies have shown that phase velocity increases for increasing age moving away from the ridge with 3.78 km/s (0- 4 Ma), 3.92 km/s (4-20 Ma), 3.98 km/s (20-50 Ma) and 4.05 km/s (50-110 Ma) at 50 s period (Forsyth et al., 1998) Another study from the GLIMPSE experiment at slightly older seafloor ages, finds that the phase velocities are higher for seafloor ages between 5-9 Ma with minimum values of 3.78 km/s at periods between 20-40 s (Weeraratne et al., 2007). Previous studies in the Pacific sea floor (ages between 20-52 Ma) indicated that Rayleigh wave phase velocities increase as a function of age and range between 3.94 - 4.16 km/s for a period range of 20-125 s (Nishimura and Forsyth, 1988) in this age bin. My study area covers an oceanic seafloor region characterized by complex breakup and fracture of the Pacific plate near shore where at least one fracture zone is found offshore. The Monterey, Morro, Arguello and Patton are identified as pseudo faults (Atwater, 1989) indicated by disruption of magnetic anomalies. The study area covers oceanic seafloor far to the west that is demonstrates uniform undisturbed magnetic anomalies oriented in a N-S direction parallel to the EPR before it subducted beneath North America, suggesting that lithospheric growth should follow the conductive cooling 12 models. The overall range of ages for the ALBACORE offshore study area is between 18-32 Ma with an unknown lithospheric thickness. The GLIMPSE study (Weeraratne et al., 2007) was carried out on young sea-floor (5-9 Ma.) west of the East Pacific Rise using Rayleigh wave dispersion suggested that the area has a thicker lithosphere of ~40 km than the younger lithosphere observed near the EPR (~20 km) indicating lithospheric growth with increasing sea-floor age. However modeling the conductive cooling predictions showed that this thickness is thinner than predicted for the half-space cooling model (Harmon et al., 2009) and other sources for volcanism and underplating may be at work in this region of the south Pacific. A recent study of P to S conversions with ALBACORE data from OBS's (Reeves et al., 2014) observed the depth of the LAB at 58 km depth west of the Patton Escarpment in ~18 Ma age seafloor. The Murray Fracture zone extends from Point Conception west ward across our study area (Figure 1.3) and separates two distinctly different areas of the seafloor (Huene, 1968). The fracture zone trends east-northeast for 1,900 miles (3,000 km) from latitude 28° N, longitude 155° W (north of the Hawaiian Islands) to Point Conception on the west coast of southern California. The maximum relief of the feature is about 6,600 feet (2,000 meters). The zone has an irregular topography of ridges, scarps, and elongate depressions (Figure 2.1). Regional depths of the seafloor north of the fracture zone are several hundred meters greater than those to the south. The patterns of magnetic anomalies area appear to be displaced laterally by 90 to 420 miles (145 to 675 km) across the fracture zone offset, and rocks of the northern block are tens of millions of years older than adjacent rocks south of the fracture zone. This eastward displacement of the seafloor north of the fracture zone is only apparent, resulting from seafloor spreading at a mid- 13 ocean ridge that was active from about 80 to 10 million years ago. The fracture zone can be dated by means of paleogeographic reconstruction of the late Cretaceous-Early Tertiary coastal sequence which suggests that the formation of the Murray Fracture Zone was later than Late Cretaceous time, but not later than middle Eocene time (Yeats, 1968). The fracture zone is mainly characterized by an age offset which is observed from the magnetic lineation’s (Figure 1.3) (Atwater, 1989). Areas of thicker lithosphere or colder material are expected to be found on the north side (older region) of the Murray Fracture zone, whereas areas of thinner lithosphere are expected to be found on the south side (younger region). 125 W 120W 14 Figure 1.3: Seafloor magnetic isochron map (Atwater, 1989). Black box shows the study area for our marine seismic deployment. 1.4 Previous Rayleigh wave study using ALBACORE data Preliminary Rayleigh wave analysis for the ALBACORE project was previously done using marine and land based data filtered from 40 s – 78 s using 39 earthquake events and 104 stations. Phase velocities were observed to be 1.3% lower than the previous studies (Nishimura and Forsyth, 1988) in 20-52 Ma seafloor for periods above 40 s . Rayleigh wave tomography results showed that the Sierra Nevada range has a very deep structural root to at least 100 km depth. The azimuthal anisotropy averaged over the study area indicates fast directions which are parallel to Pacific plate motion, N 78.5˚ W, while fast directions in the inner Borderland demonstrate a change to N-S alignment at periods longer than 40 s. Higher velocities were observed at long periods which sample the oceanic mantle can be compared to lower velocities in the continental mantle and suggested that the formation history of the lithosphere is different in the two environments. This preliminary study ignored data below 40 s in the OBS seafloor stations due to a problem with instrument amplitudes. Thus velocity structure of the crust and lithosphere offshore was poor, although resolution of the sub lithosphere was good. This study was used a limited number of events which I will add to here. 15 Chapter 2 Seismic Data 2.1 OBS Deployment and Retrieval In order to collect long term teleseismic offshore seismic data, 34 ocean bottom seismometers (OBSs) were deployed on August 2010 which recorded data until September 2011. This OBS seismic array consisted of 24 long period (LP) OBSs and 10 short period (SP) OBSs which were deployed in an array 150 km north-south by 400 km east-west off the coast of southern California (Figure 2.1). SAF NAm PAC Figure 2.1: ALBACORE OBS deployment area with bathymetry compiled by Shintaku et al., 2010. Ship tracks shown for Sept. 7-16, 2011 recovery cruise on the R/V New Horizon. Bathymetry is a compilation of ship track data sets from the NGDC, USGS, 2010 ALBACORE deployment cruise (Shintaku et al., 2010), and global data (Smith and Sandwell, 1997). Circles indicate stations with limited or no seismometer data. Coastlines are outlined in black and gray including the Channel Islands. Red arrows represent the absolute plate motion vectors. Red line on the NE quadrant marks the approximate location of the SAF. Triangles and hexagons represent long period and short period instruments respectively. Numbers correspond to deployment order and that is same as stations number. 16 This ALBACORE experiment consisted of three types of instruments, 21 long-period (LP) Trillium T-240 sensors, 3 long-period (LP) Trillium T-40 sensors, and 10 shortperiod (SP) Sercel L-28 sensors. All LP instruments were equipped with differential pressure gauges (DPGs) and all SP's included a hydrophone. All the OBS stations were set to record at a sample rate of 50 samples per second (sps) (Kohler, 2010). The cruise for the deployment phase took place on the R/V Melville from the San Diego port. The OBS locations were chosen such that the station spacing was approximately uniform, including stations deployed on the Channel Islands of the California Integrated Seismic Network Stations. In shallow-water of the Borderland region, station spacing was approximately 50 km, and on deep water seafloor, the station spacing was approximately 75 km. The depth range for the OBSs was between 1000 and 4500 m which was designed deliberately to avoid biological and sediment interference which often occurs in shallower waters. This also reduces the risk of instrument loss and damage from industrial ship activity as well as noise from shallow-water currents. On September 2011, the R/V New Horizon departed out from San Diego port in order to recover OBS instruments and the team recovered 32 instruments out of 34. Station OBS 14 was never recovered due to problem with communication that may have been due to damage of the glass flotation balls during deep water deployment (4374 m). OBS 04 communicated well, but was been regained from the seafloor possibly due to sediment burial or other obstruction. However, out of 32 recovered stations, only 26 stations had useful data. OBS 25, 16, 12 and 9 did not record any sensor data probably due to connection problems in between sensors and recording devices. Station OBS 17 17 recorded the first 3 month s of data only. So, 76% total data return was obtained from this marine deployment. 2.2 Land Stations: In this study, OBS stations were combined with land stations in western California and on the Channel Islands. Land and OBS stations recorded data for 11-12 month simultaneously. I selected 82 land stations from the California Integrated Seismic Network (CISN) between -116˚ to -125˚ longitude and 32.5˚ to 37.5˚ latitude seven of which were located on the Channel Islands during our experiment. Stations used from the CISN network consisted of 5 different high gain broad-band seismometers including 5 STS-1 (0.0027-10 Hz) seismometers, 3 STS-2.5 (0.0083-50 Hz) seismometers, 55 STS-2 (0.0083-50 Hz) seismometers, 7 CMG-3ESP (0.0083-50 Hz) seismometers and 12 CMG3T (0.0083-50 Hz) seismometers. Station spacing for the selected land stations is approximately 35 km. However, near the coast, station spacing was reduced to 20 km to utilize all operating stations along the coast. 2.3 Data Correction: A correction for instrument responses was applied to all data sets. This is an important element for the Rayleigh wave inversion procedure especially when OBS and land stations are combined with many different sensor types. Rayleigh wave analysis incorporates amplitudes in the solution for velocity, thus amplitudes must be comparable between stations. All short period OBS (SP) were significantly larger (200%) than the long period OBS (LP) (Escobar et al., 2013). So a second constant amplitude correction was applied to rectify this by determining an empirical fit comparing a group of SP data to a group of LP data (Escobar et al., 2013). 18 Then all OBS and land seismometers were corrected with respect to the response of a common reliable STS-2 instrument from the California Integrated Seismic Network (CISN). After this step, all OBS amplitudes were observed to be smaller by two orders of magnitude in amplitude compared to all land stations. Thus final constant amplitude corrections were applied to the OBS data by following the procedure mentioned above. A second correction of 0.00063 was made to the SP OBSs since the amplitudes were still offset compared with the amplitudes obtained after the first correction. Data was initially analyzed by windowing and filtering Rayleigh waves at periods from16 s to 78 s in the initial data set. A preliminary inversion was performed to solve for any additional amplitude corrections. The long period OBS's located on deep seafloor showed a frequency-dependent amplitude problem for periods below 40 s (Figure-2.2) (Escobar et al., 2013). Since this does not occur for stations located in shallow water in the Borderland, we suggest that this is not due to problems with the station or the station response, but rather to significant structural changes across the study including increasing water depth, crustal thickness and lithospheric thickness and velocities. Therefore in the initial analysis all short period data below 40 s was removed from LP OBS’s (Escobar, 2013). Several land station were also omitted such as GRA, IDO, SRI, SCZ2 and TA2 and OBS stations 2 and 17 which exhibited frequency-dependent amplitudes. 19 WD;4248 m WD;3769 m m m West WD;1730 m East Figure 2.2: Behavior of the amplitudes with increasing water depth (WD) before frequencyamplitude corrections were applied. A correction of 1.0 indicates no correction is needed. This graph shows how the OBS amplitude correction decrease dramatically at short periods for stations that are located in the west or in deep ocean water (Escobar et al., 2013). In order to correct for the frequency-dependent amplitude problem I use a 5 region starting model to guide the inversion. Using a well-defined starting model that considers the details in crustal thickness changes, sediments, underplating, and the water column functions as a damping parameter in the inversion. The study region was divided into five geologic regions for the land, inner Borderland, outer Borderland, mid–seafloor (age 15-25 Ma) and deep seafloor (age 25-35 Ma) based on geology, topography, bathymetry, and magnetic anomalies (Figure 3.1). Then P-wave, S-wave velocity structure models and density structure models for each region were compiled based on results from previous studies. I obtained velocities and densities for the water column, sediment layers, upper crust, lower crust, marine fossil underplating layers and the lithosphere. Pwave, S-wave, and density values in each region are shown in Table 2.1 with respect to thickness of each layer. Ocean drilling project reports (Vedder et al., 2007) were also used to determine sediment layer thickness especially in the Borderlands and midseafloor. S-wave velocity variation with depth for each region is shown in Figure 2.3. Then a preliminary inversion was performed to predict phase velocities for each region 20 from the P-wave and S- wave velocity structure models. Then these predicted phase velocities (Figure 2.4) were given to another inversion and that inversion try to match predicted phase velocity with observed phase velocity for each grid node in our study area by using a grid search two plane wave method. Inversion solves this by averaging all the raypaths coming to each station and then it tries to keep constant wave attenuation in between nearby stations. So this method was used to correct the frequency dependent amplitude problem in our LP OBS’s data set. Water Sediment layer 1 Sediment layer 2 Upper crust Lower crust Fossil oceanic crust Deep Seafloor 0-4.345 MidSeafloor 0-3.806 outer Borderland 1.34 inner Borderland 1.34 Land 1.4501.530f 1.5f 1.450-1.530f 1.450-1.530f - Vs (km/s) Thickness(km) 1.4501.530f - - - - Vp(km/s) Vs(km/s) ρ(g/cm3) Thickness(km) - 2.55f 1.47f 1.75f 1f - - - Vp(km/s) Vs(km/s) ρ(g/cm3) Thickness(km) Vp(km/s) 2d 5.7d 4.3f 2.49f 2.25f 2d 5.6d 3.29d 3.29d ρ(g/cm3) Thickness(km) Vp(km/s) 2.76d 4d 6.7d 2.76d 4d 6.7d Vs(km/s) 3.87d 3.87d Ρ(g/cm3) Thickness(km) 2.96d - 2.96d - 7.5a, e 5.57-5.64 (5.6) d, e 3.22-3.26 (3.24) d 2.73d 13.5a, e 6.17-6.25 (6.21) d 3.57-3.61 (3.59) d 2.85d 6d 10b 6.08d Vs(km/s) 7.5d 5.57-5.64 (5.6) d 3.22-3.26 (3.24) d 2.73d 8d 6.17-6.25 (6.21) d 3.57-3.61 (3.59) d 2.85d 6d Vp(km/s) Vs(km/s) - - 6.7d 3.87d 6.7d 3.87d - Avg. water depth(km) Vp (km/s) 21 - 3.51d 2.67d 25b 6.3d 3.64d 2.81d - Ρ(g/cm3) 2.98d 2.98d Depth to Moho (km) 10 12.3 22-27 22-27 21-37 d d d c Lithosphere Thickness (km) 60 50 68.5 23 60b Vp(km/s) 7.8d 7.8d 7.80d 7.80d 7.90d d d d d Vs(km/s) 4.51 4.51 4.51 4.51 4.57d 3 d d d d ρ(g/cm ) 3.30 3.30 3.30 3.30 3.30d Table 2.1: P-wave, S-wave, and density value variation in each region with respect to thickness of each layer. Average values are shown with in bracket (BOLD numbers). (a) ten Brink et al., 2000, (b) Yang and Forsyth, 2006, (c) Lekic and Fisher, 2011, (d) Taniya et al., 2007, (e) Brandon et al., 2013, (f) Vedder et al., 2007) Figure 2.3: S-wave velocity variation with depth for each region. Yellow diamond (Land), white square (inner Borderland), white triangle (outer Borderland), Green diamond (mid-seafloor) and blue circle (deep seafloor). 22 Figure 2.4: Phase velocity starting model shown as a function of period for each region. Regions are shown for Land (yellow diamonds), inner Borderland (white squares), outer Borderland (black triangle), Green diamond represents mid-age (18 to 25Ma) seafloor and blue circle represent deep seafloor (25 to32 Ma). The above method corrected the frequency-dependent amplitude problem below 40s for LP OBSs (Figure 2.5). Also above amplitude correction fixed amplitude problem in some of the land stations that were not used in Escobar et al., (2013) study. So I added those stations into this study. This amplitude correction gave much better results mainly because starting models include variations in thickness of the water, sediment, upper crust, lower crust, marine fossil layers and lithosphere strongly affecting to amplitude variations across study regions. However, I removed long period data (> 33 s) from SP OBS’s, due to poor signal-to-noise ratio (SNR). 23 . West East Figure 2.5: Behavior of the amplitudes with increasing depth after the frequency-dependent amplitude correction was applied. A correction of 1.0 indicates no correction is needed. This graph shows the OBS amplitudes after the amplitude correction (compared to the pre-corrected stations shown in Figure-2.2). WD is water depth. Some of the waveforms were showing notable different in the Rayleigh wave arriving as they cross the Borderlands compared to observations at more distant instruments in deeper water (see Appendix E Figure E1). This is likely due to scattering and multipathing through the sharp jumps in crustal structure between the continent, Borderland, and deep sea floor. So those events were separated as a short period land (<40 s), short period ocean ( < 40 s) and long periods land and ocean (>40 s). The SNCC land station required a large constant correction of 3.0 to be consistent with land and other instruments. A group of the long period OBS stations (OBS 010, OBS 011, OBS 013, OBS 018, OBS 022, OBS 023, OBS 028) required constant corrections of 0.3. Another group of stations (OBS 003, OBS 006, OBS 015, OBS 020, OBS 032) needed a correction of 0.5 and other OBS stations needed a constant correction that ranged from0.25 – 0.6. All short period OBS stations required 0.6 - 2.5 constant amplitude corrections except OBS 002, OBS 026, OBS 029, and OBS 031and OBS 033 which needed no correction. 24 Land stations such as GRA, IDO, JVA, PMD, PDU, CWC, CLC, CCC, RRX, MSC required a constant correction of a 1.2 and station SLR had constant correction of 1.3. I applied additional constant amplitude corrections of 0.4 -0.5 to these land stations: SMW, FIG, SCI2, LGU, SMI, NJQ, SYP, GOR, GATR, STC, SMM, BCW, PHL, MPP, MUR. See Appendix C for a full list of station corrections. 2.4 Seismic events For this study, 80 teleseismic events were used (added 41 events to the preliminary study (Escobar et al., 2013)). The azimuthal distribution of events is shown in an equidistant plot in Figure 2.6. This data set had good azimuthal distribution except in the NE azimuth due to lack of seismicity or major plate boundaries in the NE region of our study in central North America. The events with back azimuth ranging from 180˚ to 360˚ had an higher average moment magnitude compared with other azimuth. I used only the vertical component of the seismogram for Rayleigh wave energy. Rayleigh waves in this data set had good signal to noise ratio (> 3). Rayleigh wave phase velocities were analyzed at 14 different periods with a central frequency ranging between 16 – 78 seconds. Each seismogram was filtered and windowed to isolate the Rayleigh wave at each period. 25 Figure 2.6: Azimuthal distribution of 80 earthquakes used in this study projected in an equidistant plot with my study area (red square) plotted at the center. The straight black lines represent the great circle ray paths followed by each earthquake (black filled circle) from the source to the center of the seismic array. The ray path coverage for study area is shown in Figure 2.7. A high density of crossing raypaths is observed in the study area. Two quadrants lack ray path coverage in the area (NE and SW of the study area) due to a poor seismicity on the continental side and lack or high quality events from the SW Pacific. The events which originate from a SE backazimuth (raypaths parallel to coast line) show notable different in the waveform as they cross land, Borderland and seafloor (See appendix E). This is likely due to scattering and multipathing through the sharp jumps in velocity structure between the continent, Borderland, and seafloor. So for this study those events data were separated as short periods land, short periods sea and long periods. 26 Figure 2.7: Ray paths coverage for the study area considering an optimal 100 events. Only 80 events are used in the data set presented here but have an even azimuthal distribution so the pattern of coverage is the same with slightly lower density. The black lines are ray paths from the source to the station. Red triangles mark the location of the seismic stations which include OBS's and land stations. The coastline is shown in white lines. (Two plots were required to plot the full 20,000 ray paths.) 27 Chapter 3 Surface Wave Methods The phase velocity measurements obtained from Rayleigh wave paths crossing oceanic seafloor, the coastline, and the continental region is a powerful approach to detect variations in crustal, lithospheric, and asthenospheric structure. Most events exhibited variations in amplitude or wave form across the array that are indicative of focusing or multipath propagation between the source and the array caused by lateral heterogeneities (Friedrich et al., 1994). The inversions were used for this analysis taking into account perturbations in the wave field due to non-great circle path propagation with a two-plane wave approximation (Forsyth and Li., 2005). The advantage of this method compared with other method (single plane wave approximation) is that this provides 30% - 40% reduction in phase or travel time misfits (Li et al., 2003) and this approximation and allows for complexity in the incoming wave field not considered in standard great circle path tomographic techniques. Finite-frequency effects are also considered in this method (Yang and Forsyth, 2006b). This inversion uses a large number of crossing ray paths with statistical analysis to obtain velocity variation laterally and vertically across the array. The wave field for two incoming interfering plane wave represented by; (1) Where, A is the amplitude, k is the wavenumber, t the time and x is the position vector for each wave. Thus the incoming wave field at frequency ω is described by six parameters including the amplitude, phase and direction of each of the two waves. Inversions which solve for uniform phase velocity, or velocity in grouped regions use a grid search method 28 to solve for wave parameters. The inversion solves for the phase, amplitude, and propagation direction for both of the two plane waves that represent each event, as well as the averaged phase velocity (1-D) for each region (Land, Borderland and Ocean), regional phase velocity (2-D), and parameters for azimuthal anisotropy. The surface wave phase velocity is given by, (2) Where, ω is the angular frequency, θ is the azimuth, Ao is the azimuthally averaged phase velocity and A1, A2, are azimuthal anisotropic coefficients in which we have neglected terms of higher order (Smith and Dahlen, 1973; Weeraratne et al., 2007). In the data analysis, each seismogram was decimated to a sample rate of 5 sps and then filtered with a 10-mHz-wide, zero-phase-shift band pass filter centered at the frequencies form 16s -78s. Then each seismogram was windowed around the Rayleigh wave to avoid noise and other body wave phases. Next, the phase and amplitude of each seismogram were determined using Fourier analysis. The parameters of each incoming wave including phase, amplitude and direction, were solved using both a non-linear inversion (1) and a linear inversion (2). In this study, I used five different staring models (Figure 2.4) for each specific region at a given period. The inversion produces the phase velocity that gives the best fit between observed and predicted amplitude and phase for each regions. And finally, phase velocity coefficients are obtained at grid nodes spaced 0.5o apart in longitude and latitude throughout the study area (Figure 3.1). I solved for velocity using a linear inversion (2) that applies a Gaussian weighted average of the coefficients of the neighboring points. (I corrected the magenta diamonds, green circles 29 and blue circles to be consistent with the seafloor age division at 25 Ma (see Figure 1.3) between anomalies 6a and 7 (rather than using boundaries that are parallel with the coastline). Figure 3.1: Grid node configuration used to solve for phase velocities within the study area. Brown triangles mark locations of the OBS's and land stations used in this study. Red line marks location of the Murray fracture zone. Blue line marks location of the San Andreas Fault. Green line marks location of the Garlock fault. North side of the Murray fracture zone is designated by blue circles, South side of the Murray fracture zone is designated by green circles, mid ocean floor is designated by Megenta diamond, outer Borderland is designated by black triangle, inner Borderland is designated by white squares, north Garlock is designated by light blue squares, south Garlock is designated by red diamonds and west San Andreas is designated by yellow diamonds. In the case of solutions for 3 regions the north, south of the Murray fracture zones and mid ocean floor (blue circles, green circle and magenta diamonds) are grouped together as the 3 rd group, the inner and outer Borderland (black triangle and white squares) are the 2nd group and the north, south Garlock (light blue squares and red diamonds) and west San Andreas (yellow diamonds) are the 1st group. 30 Chapter 4 Results 4.1 Regional Phase Velocity results Several inversion steps were used to determine Rayleigh wave phase velocities for the study area. Previous studies solved for an average phase velocity for the entire study area (Escobar at el., 2013). I take this a step further and solve for phase velocities in 3 regions for the land, Borderland, and deep seafloor (Figure 3.1). In each region, the average phase velocity is obtained and reported as a single velocity for each period. I start with this step because the 3 regions in my study area are drastically different in crustal and lithospheric structure and averaging over the entire study area will only have limited meaning. I will perform a 3 step inversion process and solve for velocities in 3 regions in step 1, I will solve for 5 regions in step 2, and I will solve for 8 regions in step 3. These regions are chosen based on geology, topography, bathymetry, structural data, and seafloor magnetic anomalies. All inversions will solve simultaneously for anisotropy in each region where velocity is grouped. A step-by-step inversion procedure stabilizes the final phase velocity and anisotropic results. I use a starting model which has a phase velocity curve for each of 5 regions (see Figure 2.4). The starting models are based on previous independent studies from active source seismology, surface waves, and bore holes studies (see Table 2.1) and therefore contain a priori information. For the inversion, phase velocities for five regions were input (Figure 2.4) as a staring model, but the first set of inversions solved for average phase velocities for 3 regions (Figure 4.1) as output. I selected respective grid nodes (see Figure 3.1) which lie in each region by using boundary points surrounding each region. 31 Figure 4.1: Average phase velocity as a function of period in three regions. The deep seafloor is designated by blue circles, the Borderland is designated by white squares, and the land region is designated by yellow diamonds. Vertical error bars indicate two standard deviations at 95% confidence and some of them are smaller than the symbols. Phase velocities from previous studies are shown for seafloor age 20 - 52 Ma by the x symbol (Nishimura and Forsyth, 1989). Previous Rayleigh wave studies (Yang and Forsyth, 2006) for southern California using land and Channel Island stations are designated by white circles. Phase velocities are solved for simultaneously with anisotropy averaged in 3 regions. Phase velocities are higher in the ocean region (blue circles) compared to all other regions for periods between 20 s – 70 s. Velocities in the deep sea floor are 1.6% lower than previous studies in 20- 52 Ma seafloor (Nishimura and Forsyth,1988) but are 2% – 10% higher than observed previously (Yang and Forsyth, 2006) on land (white circles). Above 70 s, the difference between phase velocities beneath the seafloor and land become smaller and are not resolved within error. The two lowest points for the seafloor at 16 -18 s may be caused by averaging the entire seafloor in our study. It may also be 32 caused by sampling of the water column in deep water at short periods. I will test each possibility here. My starting model (Figure 2.4) shows the oldest seafloor and the midseafloor have very different velocities of 3.90 km/s and 3.69 km/s, at 16s and 3.95 km/s and 3.72 km/s at 18s respectively. I will test for this by dividing the seafloor study area into 2 age groups in the next section. However, I note that the average of these values is still higher than the results show in Figure 4.1. This indicates another source for the extreme low velocities may be present. The depth of the seafloor in my study area (not including the Borderlands) extends from 2700 – 4500 m. At the deepest depths, short period Rayleigh waves sample structure near the surface in deep seafloor and may be sampling the water column which will necessarily lower the phase velocities below what is expected for crustal structure. Phase velocities are lower on land (yellow diamond) compared to all other regions for all periods up to 40 s. Above 60 s, phase velocities in the Borderland, and land are indistinguishable within error. Phase velocities for land (yellow diamond) increase steeply from 16 s to 25 s with velocities vary from 3.37 km/s to 3.69 km/s. The dispersion curve above 30 s, changes slope to gentle for periods from 30 s up to 78 s with velocities vary from 3.73 km/s to 3.92 km/s. Phase velocities for these two regions are significantly lower than velocities in the deep seafloor by 2.3% at periods between 40 s – 59 s.The shape of the dispersion curve for our region on land is similar in character to previous studies performed with Rayleigh waves using only land and Channel Island stations (Yang and Forsyth, 2006b) shown in open circles, however we provide data at shorter periods starting at 16 s which shows sampling of the continental crust. Velocities between 40 s – 60 s are slightly higher (yellow diamonds) by 0.8 % in our data set compared to 33 this previous study on land (open circles). This may be caused by better crossing raypaths across the continental margin in our study afforded by the OBS deployment which will prevent laterally smoothing. This difference in the results (on land) might have occurred since previous studies cover a large area (east of our study region) including eastern California, Nevada and Arizona which have thicker crust and also contains the BasinRange province. Phase velocities for the Borderland (white squares) increase steeply from 16 s to 25 s with velocities vary from 3.61 km/s to 3.82 km/s. The dispersion curve above 30 s, changes slope to gentle for periods from 30 s up to 78 s with velocities vary from 3.85 km/s to 3.90 km/s. Sampling of the crustal velocities are indicated by the steep slope on the dispersion curve at short periods. This slope is slightly steeper in the land region where the crust to the Borderland. The crustal velocities are significantly higher overall in the Borderland compared to the land region by 2.9%. I test the stability of our phase velocity results by further breaking up the deep seafloor into two regions based on seafloor magnetic anomalies and age. This divides the study area into a total of 5 regions. The mid-seafloor region is designated at magnetic anomaly #6a (see Figure 1.3) to the Patten escarpment or coastline sampling seafloor ages between 15 - 25 Ma. The deep seafloor region samples seafloor age from 25 – 35 Ma and also exhibits more uniform magnetic anomalies oriented roughly N-S. For the inversion, phase velocities for five regions were input (Figure 2.4) as a staring model and the inversion solved for average phase velocities for these same 5 regions (land, inner Borderland, outer Borderland, mid-seafloor and deep seafloor) (see Figure 4.2). The inversion solves simultaneously for anisotropy averaged in the same 5 regions and also for 8 regions (north Garlock fault, south Garlock fault, west San Andreas fault, inner 34 Borderland, outer Borderland, mid seafloor, north Murray fracture zone and south Murray fracture zone) (see Figure 4.3) separately. Figure 4.2: Average phase velocity as a function of period in five regions. Phase velocities are solved for simultaneously with anisotropy where both parameters are allowed to vary as a group in each region. The deep seafloor is designated by blue circles, mid seafloor is designated by magenta diamond, outer Borderland is designated by black triangle, inner Borderland is designated by white squares, and land is designated by yellow diamonds. Vertical error bars indicate two standard deviations and most are smaller than the symbols. Phase velocities are higher in the mid-seafloor region (magenta diamond) by 2.5% compared to the deep seafloor region (blue circle) for periods between 18 s – 33 s. However, above 45 s phase velocities are higher in deep sea floor compared with midseafloor region. At about 16 s phase velocities in the deep seafloor and mid-seafloor are very low and are likely caused by sampling of the water column. Sampling of the water 35 column may affect velocities in the deep seafloor region to periods up to 30 s causing these velocities to appear lower than the mid-seafloor region. We also note that several OBS stations in the deepest seafloor region did not return sensor data, and I only have data from 4 stations within this group, whereas I use 9 stations in the mid-seafloor group. Future work should include DPG data (DPG data can be converted to verticals Teodor et al., 2013) from the missing sensor instruments to improve this problem. Phase velocities closely in the inner Borderland region (white squares) and the outer Borderland region (black triangle) for all periods up to 50 s. This dispersion curve has a steep slope with velocities between 3.53 km/s to 3.84 km/s between 16 s – 25 s period in the both region but it extends to at least 30 s in the outer Borderlands. Phase velocities in the inner Borderland are significantly higher than the outer Borderland at all periods above 50 s by 2.4% and are as high as the mid-seafloor region. The slope in the outer Borderlands is negative for periods above 33 s where velocities decrease from 3.87 km/s to 3.82 km/s indicating a low velocity zone. At periods above 40 s, the velocities in the outer Borderland are as low as on land, and are the lowest of all regions in our study above 55 s. The velocities on land (yellow diamonds) are similar to the previous inversion for three regions, however, above 40s, velocities on land are significantly higher than the Yang and Forsyth (2006b) result (white circles). This suggests that previous studies may have lateral averaging from offshore structure that was mapped into the land velocities which we are able to improve by having offshore stations and more crossing raypaths across the continental margin. 36 Figure 4.3: Average phase velocity as a function of period in eight regions. Phase velocities are solved for simultaneously with anisotropy where both parameters are allowed to vary as a group in each region. The north side of the Murray fracture zone is designated by blue circles, south side of the Murray fracture zone is designated by green circles mid seafloor is designated by magenta diamonds, outer Borderland is designated by black triangles, inner Borderland is designated by white squares, west San Andreas fault region is designated by yellow diamonds, North Garlock fault region is designated by light blue squares and South Garlock fault region is designated by red diamonds. Vertical error bars indicate two standard deviations and most are smaller than the symbols. In the third inversion step, I break up the continental land region into 3 smaller regions and the deep seafloor into 3 smaller regions outlined by major fault boundaries and keep the inner and outer Borderland regions separated which divides are study area into 8 total region. (see Figure 3.1). I designate north side of the Murray fracture zone with blue circle and south side of the Murray fracture zone by green circle. Also I designate the land region west of the San Andreas faults everywhere in our study area 37 with yellow diamonds, the region east of the San Andreas fault and south of the Garlock faults by red diamonds, and the region east of the San Andreas fault and north of the Garlock fault by light blue squares. The inversion obtains the average phase velocity in each region and solves simultaneously for anisotropy which is also averaged in each of the 8 regions. Phase velocity increases from 3.54 km/s to 3.79 km/s in north Murray fracture zone (blue circle) for periods between 18 s – 40 s. This is much lower than all regions except on land. This suggests sampling of the water column and averaging of paths which cross the Pacific in very deep water before they arrive at the western edge of our station coverage may affect velocities in the region north of the Murray fracture zone to periods up to 40 s causing these velocities to appear lower than the south Murray fracture zone, mid-seafloor region and Borderland regions. However, above 45 s phase velocity increase gradually up to 3.91 km/s. Phase velocity at 16 s follow starting model and it may not be controlling by the data. Phase velocity in the south Murray fracture zone increases from 3.64 km/s to 3.93 km/s for periods between 18 s- 40 s. But above 40 s, it increases gradually up to 3.92 km/s. Phase velocity at 16 s is 3.66 km/s. North and south Murray fracture zones show low phase velocity compared to mid seafloor for periods between 18 s – 33 s and are likely caused by sampling of the water column. Phase velocities are highest on land in the region south of the Garlock fault (red diamond) by 1.8% compared to the north Garlock and west San Andreas regions (light blue squares and yellow diamonds) for periods between 25 s – 45 s. However, phase velocities are same in all 3 land regions (the south Garlock, north Garlock and west San Andreas regions) for periods between 16 s – 22 s. Also phase velocities in the south Garlock fault region are higher by 1.9% compared to Yang and Forsyth (2006b). Above 38 60 s, phase velocities in all three land regions are similar except above 65 s, velocities in the north Garlock region are slightly higher. Velocities in the offshore Borderland and mid seafloor regions are not significantly affected by the division of the land and deep seafloor regions and are similar to the results in Figure 4.2. 4.2 Regional anisotropy results Seismic anisotropy was obtained in all the phase velocity inversions discussed above. Here I show anisotropy results for inversions which consider 8 regions (North Murray fracture zone, south Murray fracture zone, mid-seafloor, outer Borderland, inner Borderland, north Garlock region, south Garlock region and San Andreas region). In this inversion, a statistical average of hundreds of raypaths passing through each of the 8 regions are used to solve for the dominant azimuth that displays the fastest phase velocity and the strongest anisotropic direction in that region. Data is carefully selected and chosen by hand. Thus some periods have a slightly different number of raypaths than others even within the same region. The availability of stations as well as good quality earthquakes from every testable azimuth will also influence the final result. The statistically averaged dominant anisotropic azimuth is displayed as a function of period showing variations in anisotropic as a function of depth, where the depth of peak sensitivity is approximately 4/3 period. Azimuthal anisotropy is well resolved from zero in all regions for near all periods. In the north Murray fracture zone (Figure 4.4a), the fast direction of anisotropy is consistent at the shortest periods of 20 s – 40 s is WNW-ESE (N 80˚ W ± 0.9˚) with in error and shifts to turn NW - SE (N 44˚ W ± 4.0˚) from 50 s – 59 s periods with in error. The fast direction changes for periods above 40 s indicating a consistent fast direction of NW-SE for nearly all periods from 45 s – 59 s and, 69 s and 39 78 s which show a WNW-ESE (N 83˚ W ± 7.0˚) fast direction. The strong consistency of azimuthal direction over more than 3 consecutive periods suggests this observation is real and that a region with 2 layers of anisotropy above and below 45 s is well resolved. At 16 s and 18 s errors are greater than 10˚ and not well resolved. The magnitude of the anisotropy for 16 s - 40 s except 18 s is unrealistically high and is likely caused by unreasonably low phase velocities at these periods in the deep seafloor (Figure 4.1, 4.2 and 4.3 blue circles) which sample the water column. This exemplifies the tradeoff between phase velocity and anisotropy in our inversion (see the 3 terms in equation (2)).The magnitude of the anisotropy in north Murray fracture zone is highest at short periods and decreases from 9.4 % at 20 s down to 4.5% at 33 s within error. This decrease in the strength of the anisotropy may be due to Rayleigh waves with large wavelengths sampling two structures which have a boundary at approximately the peak sensitivity depth of 44 km depth. The extraordinarily high magnitude of anisotropy even at 20 s may indicate sampling of the water column which lowers phase velocities and artificially raises anisotropy results. The magnitude of anisotropy may be constant above 40 s at about 2.0 % within error. 40 Figure 4.4a: Anisotropy in the north Murray fracture zone area. The error bars indicate two standard deviations at 95 % confidence. Long lines indicate the direction of the anisotropy in map view with reference to north (arrow). Arrow with PAC shows Pacific plate motion direction Figure 4.4b: Anisotropy in the south Murray fracture zone area. The error bars indicate two standard deviations at 95 % confidence. Long lines indicate the direction of the anisotropy in map view with reference to north (arrow). Arrow with PAC shows Pacific plate motion direction. In the south Murray fracture zone (Figure 4.4b), the fast direction of anisotropy is NW - SE (N 65˚ W ± 8.0˚) for all periods with in error except 69 s. The fast direction is consistent with pacific plate motion direction (N 78.5˚ W, Gripp and Gordon (2002)) for periods between 20 s to 59 s except 50 s and 55 s. The magnitude of the anisotropy for 16 s and 18 s is unrealistically high by 7% (not shown in figure 4.4b) and is likely caused by unreasonably low phase velocities at these periods in the deep seafloor (Figure 4.1, 4.2 blue circles and 4.3 green circles) which sample the water column. This exemplifies the tradeoff between phase velocity and anisotropy in our inversion (see the 3 terms in equation (2)).The magnitude of the anisotropy in south Murray fracture zone is highest at short periods and decreases from 3.1 % at 20 s down to 0.3% at 40 s within error. This 41 decrease in the strength of the anisotropy may be due to Rayleigh waves with large wavelengths sampling two structures which have a boundary at approximately the peak sensitivity depth of 47 km depth. The extraordinarily high magnitude of anisotropy even at 20 s may indicate sampling of the water column which lowers phase velocities and artificially raises anisotropy results. The magnitude of the anisotropy is constant above 40 s at about 1.0% with in error. Figure 4.5: Anisotropy in the mid age seafloor area. The error bars indicate two standard deviations, 95 % confidence. Long lines indicate the direction of the strength of the anisotropy for fast directions. Arrow with PAC shows Pacific plate motion direction Anisotropy for the mid-seafloor region (Figure 4.5) indicates that the fast direction is WSW-ESE (S 80˚ W ± 2.0˚) at all short periods at 50 s and below 50 s. A transition is observed to NW-SE (N 77˚ W ± 4.0˚) fast direction at longer periods above 50 s. The strength of anisotropy is in the range of 0.2 – 2.0 % for all periods below 78 s except for shortest periods (< 25 s). The high magnitudes in anisotropy at short periods are likely to due to the indirect result of sampling of the water column. Above 55 s, 42 anisotropy increases in magnitude at 59 s and 79 s. The magnitude of the anisotropy for 16 s is unrealistically high at 10.0% and is likely caused by unusual low phase velocities at these periods in the mid-seafloor (Figure 4.2 and 4.3) which sample the water column. The anisotropy in the outer Borderland (Figure 4.6) is approximately E-W at 40 s and is remarkably consistent for periods up to 55 s. The periods from 20 s – 33 s demonstrate a slight rotation to a NE-SW (N 67˚ E ± 2.0˚) orientation. The fast direction changes to a NW-SE direction (average is N 60˚ W ± 4.0˚) for periods above 55 s except 78 s and strength of the anisotropy decreases to about 2.1% for long periods (40 s – 55 s) within error. The average fast direction in the inner Borderland (Figure 4.7) is NW-SE (average is N 77˚ W ± 5.0˚) for most periods up to 45 s (although period of 33 s indicate an E-W azimuth). Above 45 s, the average fast direction is uniform for all these periods directed to NE–SW (N 43˚ E ± 7.0˚). The decrease in magnitude at 45 s may be due to Rayleigh waves with large wavelengths which sample two layered structures which have a boundary at approximately the peak sensitivity depth of 50 km depth. The magnitude of the anisotropy for 16 s is unrealistically high at 5.4% (plots outside the graph - not shown) and is likely caused by unreasonably low phase velocities at this period in the inner Borderland (Figure 4.1, 4.2 and 4.3 white squares). The anisotropy in the inner Borderland decreases from 18 s to 45 s from 3.7% to 0.2%. Anisotropy for periods above 45 s may be constant at 1.4 % within error. 43 Figure 4.6: Anisotropy in the outer Borderland. The error bars indicate two standard deviations, 95 % confidence. Long lines indicate the direction of the strength of the anisotropy for fast directions. Arrow with PAC shows Pacific plate motion direction. Figure 4.7: Anisotropy in the inner Borderland. The error bars indicate two standard deviations, 95 % confidence. Long lines indicate the direction of the strength of the anisotropy for fast directions. Arrow with PAC shows Pacific plate motion direction. 44 Figure 4.8: Anisotropy in the north Garlock fault region. The error bars indicate two standard deviations, 95 % confidence. Long lines indicate the direction of the strength of the anisotropy for fast directions. Arrow with PAC shows Pacific plate motion direction. In the north Garlock region (Figure 4.8), the direction of anisotropy is trending NW – SE (N 55˚ W ± 7.0˚) for all periods up to 20 s. The fast direction changes at 25 s to an E-W azimuth that remains constant up to 40 s. The fast direction rotates slightly to a NW-SE (N 60˚ W ± 10.0˚) direction for periods 45 s – 60 s. Finally, at the longest periods of 67 s – 78 s, the fast direction indicates N-S orientation. The average magnitude of the anisotropy is 1.3% with the exception of anomalous points at 18 s and 33 s. 45 Figure 4.9: Anisotropy in the south Garlock fault region. The error bars indicate two standard deviations, 95 % confidence. Long lines indicate the direction of the strength of the anisotropy for fast directions. Arrow with PAC shows Pacific plate motion direction. In the south Garlock region (Figure 4.9), the direction of anisotropy is trending NW – SE and is remarkably consistent for all periods from 16 s – 78 s. The average fast direction is N 67˚ W. The magnitude of the anisotropy increases steadily from 1% at short periods to 4.0% at long periods. The direction of the anisotropy in the west San Andreas region (Figure 4.10) is NW–SE (N 67˚ W ± 10.0˚) for all periods except 55 s and 78 s. The magnitude of the anisotropy is high and fluctuates at short periods but it decreases from 4% to 0.2% for periods between 20 s – 50 s. And then strength of the anisotropy gradually increases up to 1% for all periods above 50 s. 46 Figure 4.10: Anisotropy in the west San Andreas Fault region. The error bars indicate two standard deviations, 95 % confidence. Long lines indicate the direction of the strength of the anisotropy for fast directions. Arrow with PAC shows Pacific plate motion direction. 47 Chapter 5 Discussion 5.1 Deep seafloor Phase velocities for deep seafloor region are the highest in our study compared to the Borderland and land tectonic regions for all periods above 25 s. Phase velocities at 3.95 km/s for our study area where seafloor age varies from 15-32 Ma are consistent with previous studies for a specific age bin of 20 – 52 Ma which used only rays crossing oceanic paths recorded by land stations (Nishimura and Forsyth, 1988). Below 30 s, Phase velocities in the deep seafloor are anomalously low for short periods from 16 s – 25 s which we suspect is due to sampling of the water column in deep seafloor for short period Rayleigh waves. I divide the deep seafloor in my study area into 2 regions with ages ranging from 15 – 25 Ma (mid-seafloor) and 25 – 32 Ma (deep seafloor) as shown in Figure 3.1 and Figure 4.2. Rayleigh waves traveling through the mid-seafloor should have a small influence from deep water sampling and is exemplified by phase velocities which are very high at short periods from 20 s - 44 s at about 3.9 km/s. Again, I divide the deep seafloor in my study area into 2 regions (north Murray fracture zone and south Murray fracture zone) as shown in Figure 3.1 and Figure 4.3. Rayleigh waves traveling through the south Murray fracture zone should have a small influence from deep water sampling and is exemplified by phase velocities which are very high at short periods from 20 s - 44 s at about 3.93 km/s. Phase velocities in the deep seafloor region are significantly higher than the mid-seafloor region for periods above 45 s consistent with oceanic lithosphere that is older and colder due to conductive cooling over time. 48 Anisotropy in the north Murray fracture zones region exhibits a fast direction that is WNW-ESE for periods up to 40 s that is parallel to the absolute Pacific plate motion direction. Periods above 55 s show a fast direction that is rotated to the NW-SE direction for periods from 50 s – 60 s which may occur in the asthenospheric mantle. As this region has a few missing OBS stations and less control on lateral variations across a long raypath which crosses the Pacific ocean before arriving at the edge of our array, addition of DPG data to this region may improve or perhaps even change this result. The last few periods between 69 s – 78 s indicates a change to WNW-EE orientation that is consistent with the Pacific plate motion direction, N 78.5˚ W (Gripp and Gordon, 2002). Anisotropy in the mid-seafloor has a few subtle differences from the deep seafloor. This region of our study has the best coverage from long period OBS instruments and best control on lateral variations from crossing raypaths. Anisotropic alignment at short periods below 55 s is remarkably consistent and is WSW-ESE indicating a slight rotation from the WNW-ESE observation in short periods in the north and south Murray fracture zone regions. This is consistent with the slightly rotated magnetic anomalies to the NE-SW along the arcuately shaped Arguello and Patton pseudo fractures (Figure 1.3). This indicates fossil anisotropy was frozen in during lithospheric formation (18 Ma – 32 Ma) and this region of the seafloor underwent deformation and plate capture (Atwater, 1989) after this time when the East Pacific Rise subducted beneath North America 37 Ma. The rotated anisotropic fabric that is observed up to 50 s period must extend to about 67 km depth indicating that the magnetic anomalies on the surface indicate disruption, deformation, rotation, and plate capture of the entire lithosphere. The transition in anisotropy to a WNW-ESE direction is more 49 pronounced and consistent over a wider period range of 55 s – 79 s in the mid-seafloor and indicates alignment with active Pacific plate motion today at sublithospheric depths in the asthenosphere. The observation of anisotropic fabric in a WNW-ESE direction in the lithosphere of the mid-seafloor region is also observed in SKS splitting measurements for OBS stations East of 124˚W (Ramsay et al., 2014) indicating the SKS vertically incident raypath is at least partially sampling the lithospheric mantle. 5.2 Oceanic versus continental lithosphere The highest phase velocities are in the ocean region compared to land region for periods above 20 (Figure 4.1). Phase velocities in the seafloor (blue circles) are 2% - 10% higher than observed on land (yellow diamonds and white circles). This drastic difference indicates the formation processes of the crust and lithosphere beneath land and ocean are different. Oceanic crust and lithosphere forms at the spreading center from melted magmas derived from the depleted mantle asthenosphere (e.g. Workman and Hart, 2004). The lithosphere spreads and cools with time for 15 - 32 Ma. However, continental crust and lithosphere forms due to melting processes (western U.S.) associated with oceanic plate subduction. This continental formation process makes hydrated and rich hydrous phases where seismic waves travel slower compared to ocean region that has a more depleted dry oceanic lithosphere where seismic waves travel faster. Above 60 s, phase velocities in the inner Borderland (white squares in Figure 4.2), and land are indistinguishable within error. This suggests the lithospheric mantle is similar in these 2 regions and are very different from the oceanic mantle where velocities are significantly 50 (1.3%) higher in the deep ocean region. The velocities on land (Figure 4.2 yellow diamonds) are similar to the previous inversion for three regions, however, above 40s, velocities on land are significantly higher than the Yang and Forsyth (2006b) result (white circles). This suggests that previous studies may have lateral averaging from offshore structure that was mapped into the land velocities which we are able to improve by having offshore stations and more crossing raypaths across the continental margin. Anisotropy for continental Southern California (Figure 4.8) in the block east of the San Andreas fault and south of the Garlock fault is approximately N 67˚ W and is consistent with previous studies (Yang and Forsyth, 2006a). This is roughly parallel to the axis of the bend in the San Andreas fault and is not consistent with North American plate absolute motion (S 57 ˚ W) suggesting local shear from transform motion across the fault is dominant. However, more stations in this region would improve resolution of lateral variations. This may be consistent with escape of the block south of the Garlock fault accommodated by left lateral motion above and right lateral motion below along the SAF pushing the block to the SE direction (Dolan et al., 2012). The strong increase in the strength of anisotropy with period suggests that shearing is active to lithospheric depths. The fast wave direction of anisotropy for continental region west of the San Andreas fault for all periods except 55 s and 78 s displays a NW-SE azimuth consistent with Pacific plate motion. At About 30 s,33 s, 40 s and 59 s, the fast direction shifts slightly to a WNW-ESE azimuth for nearly all periods which is similar to the azimuth of anisotropy in the shortest periods of the offshore regions of our study in the Borderlands and deep seafloor which are perpendicular to the magnetic lineation’s offshore. While lateral smearing is possible at the coastline, this does not seem to be a factor at short periods 51 below 30s, thus we suggest this WNW-ESE fabric is present at lithospheric depths below 35 km in the continent west of the San Andreas. This observation is consistent with previous shear wave splitting results (Polet and Kanamori, 2002). Anisotropy may change again at 70 s (corresponding to 90 km depth) to a NW-SE azimuth parallel to Pacific plate motion, but this is only observed for 1 or 2 periods and is not well resolved. The magnitude of the anisotropy in south Garlock region decreases in between 20 s – 22 s, then increases until 29 s and again increases up to 78 s. So this anisotropy variation may be due to Rayleigh wave sampling of two structures which have a boundary at approximately the peak sensitivity depth of 30 km depth (±20km). This result is consistent with Reeves et al., (2014) which states that average Moho depth below Los Angeles equal to 28.5 km. 5.3 The Borderlands Phase velocities are higher in the averaged Borderland regions from 16 s – 59 s compared with land (Figure 4.1). This high phase velocity may be due to crustal accretion or underplating of sequences like the Catalina schist (Bohannon and Geist, 1998) during rifting and extension of the borderland region. Samplings of the crustal velocities are indicated by the steep slope on the dispersion curve at short periods. This slope is slightly steeper in the land region where the crust to the Borderland. This is likely produced by a thicker crust on land compared to the Borderland, consistent with active sources studies in the Borderlands that give 19 km to 23 km for the crustal thickness (ten Brink 2000) compared to 40 km on land (LARSE study: Kohler and Davis, 1997; Zhu and Kanamori, 2000). The crustal velocities shown in Figure 4.1 are significantly higher overall in the Borderland compared to the land region by 6.8 %. This suggests basalt compositions and 52 crustal underplating in the Borderlands. The inner and outer Borderland are very similar for periods up to 50s, but depart significantly at longer periods where velocities in the inner Borderland are 2.2% higher. At the longest period inner Borderland velocities approach velocities of the oldest seafloor suggesting oceanic lithospheric fabric may be present in this region at depths below 65 km. There is a strong low velocity zone in the outer Borderlands for periods above 30 s corresponding to lithospheric depths which may indicate rifting or the presence of partial melt. Anisotropy for outer Borderland (Figure 4.6) is SW – NE at periods between 20 s to 33 s. This is perpendicular to the rotated magnetic anomalies in the Arguello and Patton fracture zones and indicates this deformation extends to depths down to 44 km ± 10 km. The anisotropy is E – W at periods between 33 s to 55 s. This is may indicate the original fabric from N-S magnetic anomalies before rotation occurred. Another transition in anisotropy is observed in the outer Borderland to a NW-SE fast direction at longer periods above 55 s (and deeper depths). This suggests alignment due to mantle drag from absolute plate motion in the NW direction (Liu and Bird, 2002) may be active in the asthenosphere. This change in anisotropic fabric will define base of the lithosphere in shear wave inversions. Anisotropy in the inner Borderland is NW-SE which is rotated slightly CW from the PAC motion direction and may signify influence from CW rotation of the Transverse Range system. Anisotropy is distinctly different for all periods above 50 s in the inner Borderland with an NE-SW azimuth. This azimuth is perpendicular to the coast line and the direction of subduction before the inception of the San Andreas fault. The anomalously high phase velocities in the inner Borderland which are higher than the outer Borderland at all periods above 50 s by 2.2% and are as high as the mid53 seafloor region suggests a high velocity structure may be present. This feature may be over 65 - 80 km thick. This result is consistent with Lekic et al., (2011)which states that lithospheric thickness beneath inner Borderland is approximately 50 km and an anomalously high velocity structure is imaged beneath this region dipping to the NE (Reeves et al., 2014). 54 Chapter 6 Conclusion The anisotropic phase velocities in ocean region show 1.6% lower compared to Nishimura and Forsyth (1988). However, our result is consistent with this study. And this difference may be due to different age bin. Also our phase velocities (in land) are 0.8% higher compared to Yang and Forsyth (2006b) but our results for land and Borderland are consistent with this study. Phase velocities in the seafloor are 2% - 10% higher than observed on land. This may be caused because formation processes help to form different crust and lithosphere beneath land and ocean. Also Borderland has higher phase velocities than land but lower phase velocity than ocean may be caused because formation processes help to form different crust and lithosphere beneath land, Borderland and ocean. Anisotropy for continental Southern California is consistent with previous studies (Yang and Forsyth, 2006a). And fast wave direction of the anisotropy for land and inner Borderland are consistent with previous shear wave splitting results (Polet and Kanamori, 2002). Anisotropy for outer Borderland is consistent with Pacific plate motion, N 78.5˚ W (Gripp and Gordon, 2002). 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Res., 105, 2969-2980, 2000. 60 Appendix A: Amplitude Correction Plots Ocean Bottom Seismometer (OBS) 61 62 63 64 Land Stations 65 66 67 68 69 70 71 72 Appendix B: Earthquakes DATE O-TIME LAT LON DPT MAG DIST DLTA BAZ QLTY 2010 08 27 192349.5 35.49 54.47 7 5.7 12350 111.3 3.5 2010 09 03 111606.6 51.45 -175.87 23 6.5 4866 43.8 311.5 1 2010 09 03 163547.8 -43.52 171.83 12 7 10964 98.7 223.1 2 2010 09 07 161332.2 -15.88 -179.31 10 6.3 8318 74.9 239 3 2010 09 08 113731.9 -20.67 169.82 10 6.3 9538 85.8 242 3 2010 09 09 72801.72 -37.03 -73.41 16 6.2 9169 82.5 144.1 3 2010 09 13 71549.63 -14.61 -70.78 179 5.9 7445 67 127.2 3(SP) 2010 09 17 192115 36.44 70.77 220 6.3 12190 109.7 350.8 3 2010 09 22 80014.32 -13.39 -76.07 50 5.7 6962 62.7 130.5 3 2010 09 27 112246.1 29.65 51.67 26 5.8 12971 116.9 6.7 2010 09 29 171125.9 -4.96 133.76 26 7 11807 106.3 274.9 2 2010 10 04 132838.9 24.27 125.15 32 6.3 10612 95.5 303.7 2 2010 10 08 32613.71 51.37 -175.36 19 6.4 4830 43.5 311.4 2 2010 10 09 15404.6 10.21 -84.29 91 5.8 4458 40.1 116.9 2 2010 10 30 151833.1 -56.59 -142.29 10 6.4 10189 91.7 192.1 2 2010 11 12 190130 -35.96 -102.21 10 5.8 7896 71.1 164.8 2 2010 11 30 154559 39.8 -71.93 6 3.9 4423 39.8 66.2 5 2010 12 02 31209.82 -6 149.98 33 6.6 10366 93.3 265 1.5 2010 12 05 214435.8 -36.23 -100.83 12 5.9 7960 71.6 163.8 2 2010 12 08 52435.26 -56.41 -25.74 29 6.3 13269 119.4 140.5 3.5 2010 12 20 184159.2 28.41 59.18 12 6.7 13145 118.4 359.3 3 2010 12 21 171940.7 26.9 143.7 14 7.4 8943 80.5 295.8 2.5 2010 12 22 214938.8 26.81 143.6 10 6.4 8957 80.6 295.8 2 2010 12 23 140032.3 53.13 171.16 18 6.3 5742 51.7 314.2 2 73 2.5 2.5(SP) 2010 12 25 131637 -19.7 167.95 16 7.3 9627 86.6 243.9 1 2010 12 29 65419.64 -19.66 168.14 16 6.4 9608 86.5 243.8 1 2010 12 30 125522 40.43 -85.91 5 3.8 3232 29.1 65.9 5 2011 01 02 202017.7 -38.37 -73.35 24 7.1 9290 83.6 144.9 3(SP) 2011 01 09 100344.3 -19.16 168.31 24 6.6 9560 86 244.1 2(SP) 2011 01 13 161641.6 -20.63 168.47 9 7 9646 86.8 242.8 1 2011 01 18 202323.5 28.78 63.94 68 7.2 13106 117.9 356.1 3(SP) 2011 01 27 83828.66 28.19 59.01 12 6.2 13169 118.6 359.5 3 2011 01 29 65526.13 70.94 -6.68 6 6.2 7310 65.8 19.3 3 2011 02 04 135344.2 24.64 94.68 66 6.3 12604 113.4 325.6 3(SP) 2011 02 11 200530.8 -36.47 -73.12 27 6.8 9136 82.2 143.6 3 2011 02 12 25315.1 0.08 -17.02 10 5.6 11198 100.8 82.7 3 2011 02 14 34009.92 -35.38 -72.83 21 6.6 9056 81.5 142.7 3 2011 02 18 231532 30.08 -88 5 3.5 3175 28.6 87.7 5 2011 03 01 5346.34 -29.7 -111.98 10 6.1 7034 63.3 172.2 1 2011 03 02 185049.4 8.59 -76.86 44 5.8 5201 46.8 111.9 1 2011 03 06 123159.7 -18.04 -69.34 118 6.3 7825 70.4 128.7 3(SP) 2011 03 06 143236.1 -56.42 -27.06 87 6.5 13191 118.7 140.7 3 2011 03 07 936.45 -10.35 160.77 22 6.4 9646 86.8 255.5 5 2011 03 09 24520.33 38.44 142.84 32 7.3 8337 75 306.2 1 2011 03 09 212402.7 38.29 142.8 22 6.4 8349 75.1 306.1 2 2011 03 11 1448.33 -53.21 -117.84 10 5.7 9596 86.4 178.7 5 2011 03 11 54624.12 38.3 142.37 29 9 8380 75.4 306.3 1 2011 03 11 61540.92 36.27 141.11 48 7.9 8594 77.3 305.1 3 2011 03 11 81924.38 36.17 141.56 6 6.5 8565 77.1 304.8 5 2011 03 12 14715.4 37.59 142.65 20 6.5 8400 75.6 305.6 3 2011 03 17 24800.03 -17.27 167.83 17 6.3 9476 85.3 245.9 2 74 2011 03 22 71845.38 37.24 144 11 6.5 8317 74.9 304.7 2 2011 03 22 91906.23 37.33 141.79 31 6.3 8480 76.3 305.7 2 2011 03 22 133128.5 -33.1 -15.98 10 5.9 13104 117.9 112.8 1 2011 03 24 135511.9 20.69 99.84 8 6.8 12675 114.1 318.9 3 2011 03 29 160435.6 46.57 -76.96 18 3.5 4000 36 55.3 5 2011 03 31 1158.3 -16.54 -177.52 15 6.4 8226 74 237.3 5 2011 04 01 132910.7 35.66 26.56 59 6.1 11603 104.4 27.6 3 2011 04 07 143243.3 38.28 141.59 42 7.1 8440 76 306.6 2 2011 04 07 204151.5 17.11 -85.08 7 5.9 3926 35.3 108.8 2 2011 04 11 81612.73 37 140.4 11 6.6 8604 77.4 306 2 2011 04 18 130302.7 -34.34 179.87 86 6.6 9771 87.9 225.9 2 2011 04 18 153803.9 32.46 -65.6 10 4.6 5151 46.4 75.1 5 2011 04 23 41654.72 -10.38 161.2 79 6.8 9609 86.5 255.2 2 2011 04 30 81916.29 6.88 -82.35 8 6.2 4862 43.8 118.7 2 2011 05 10 85509.05 -20.25 168.25 11 6.9 9639 86.8 243.3 2 2011 05 13 224756.7 10.06 -84.31 89 6 4467 40.2 117.1 3 2011 05 15 130813.1 0.57 -25.65 10 6.1 10370 93.3 87.1 2 2011 05 15 183710.4 -6.13 154.41 40 6.5 9964 89.7 262.5 3 2011 05 21 1625.56 -56.07 -27.11 48 5.9 13176 118.6 140.4 5 2011 06 01 125522.4 -37.58 -73.69 21 6.3 9201 82.8 144.6 3 2011 06 05 115112 -55.84 146.62 3 6.4 13198 118.8 220 2 2011 06 13 143123 2.52 126.46 61 6.3 12004 108 285.6 4 2011 06 16 335.79 -5.93 151.04 16 6.4 10263 92.4 264.5 5 2011 06 16 72040.18 48.27 -69.65 23 3.8 4561 41.1 53.1 5 2011 06 20 163601.2 -21.7 -68.23 128 6.4 8200 73.8 130.5 3 2011 06 22 215052.3 39.96 142.21 33 6.7 8298 74.7 307.9 1.5 2011 06 24 30939.48 52.07 -171.84 52 7.3 4592 41.3 312.7 1 75 2011 06 26 121638.6 -2.38 136.63 17 6.3 11385 102.5 275.4 2 2011 07 06 190318.3 -29.54 -176.34 17 7.6 9143 82.3 227 2 2011 07 10 5710.8 38.03 143.26 23 7 8328 75 305.7 2 2011 07 11 204704.3 9.51 122.17 19 6.4 11905 107.1 294.1 2.5 2011 07 23 43424.18 38.9 141.82 41 6.3 8387 75.5 307.1 2.5 2011 07 24 185124.5 37.73 141.39 35 6.3 8487 76.4 306.2 1 2011 07 25 5051.29 -3.2 150.59 35 6.3 10140 91.3 267 3 2011 07 27 230029.8 10.78 -43.38 7 5.9 8113 73 88.2 1 2011 07 30 185351 36.94 141.14 36 6.3 8552 77 305.7 2.5 2011 07 31 233856.6 -3.52 144.83 10 6.6 10693 96.2 269.9 3 2011 08 10 234542.5 -7.18 -12.66 9 6 12042 108.4 86.5 1 2011 08 14 12939.34 -1.34 -14.65 10 5.6 11503 103.5 82.6 1 2011 08 20 165502.8 -18.36 168.1 32 7.2 9525 85.7 244.9 3 2011 08 20 171309.3 -18.33 168.11 60 6.5 9522 85.7 244.9 2 2011 08 20 181923.5 -18.31 168.22 28 7.1 9511 85.6 244.8 3 2011 08 23 175104.6 37.94 -77.93 6 5.8 3803 34.2 70.3 1 2011 08 24 174611.5 -7.64 -74.51 145 7 6607 59.5 124.8 2011 08 25 50750 37.94 -77.9 5 3.9 3933 35.4 70.1 2 2011 09 02 105554.1 52.21 -171.74 35 6.9 4587 41.3 312.9 1 2011 09 03 44857.3 -56.43 -26.83 84 6.4 13205 118.8 140.7 3 2011 09 03 225539.6 -20.65 169.75 171 7 9542 85.9 242.1 3 2(SP) O-TIME: Origin time, LAT: Latitude, LON: Longitude, DPT: Depth, MAG: Magnitude, DIST: Distance, DLTA: Delta, BAZ: Back azimuth, QLTY: Quality (1- high, 2- good, 3-middle, 4-bad, 5- very bad, SPSeparated) 76 Appendix C: Station Correction Station Corr. Station Corr. Station Corr. Station Corr. obs002 1 obs026 1 SES 0.6 CLC 1.2 obs003 0.5 obs027 2.5 MPI 0.6 CCC 1.2 obs005 1.2 obs028 0.3 MSC 1.2 LGU 0.5 obs006 0.5 obs029 1 PMD 1.2 DEC 0.8 obs007 0.25 obs030 1.3 GATR 0.5 RRX 1.2 obs008 0.25 obs031 1 SMM 0.5 SMI 0.5 obs010 0.3 obs032 0.5 FIG 0.5 TFT 0.6 obs011 0.3 obs033 1 STC 0.5 SLR 1.3 obs013 0.3 obs034 0.6 FMP 0.6 BAK 0.7 obs015 0.5 SMW 0.5 LCP 0.5 NJQ 0.5 obs017 1 LGB 0.7 SPF 0.6 SCI2 0.4 obs018 0.3 IDO 1.2 PHL 0.5 SNCC 3 obs019 0.4 BCW 0.5 PDU 1.2 SBI 0.5 obs020 0.5 SBB2 1.1 USC 0.8 SYP 0.5 obs021 1 SBC 0.6 CIA 0.6 MPM 1.2 obs022 0.3 JVA 1.2 GRA 1.2 SCZ2 0.7 obs023 0.3 CWC 1.2 GOR 0.8 obs024 0.6 WGR 0.5 MPP 0.5 Note: stations are only listed for which corrections are made. All other stations have a correction value of 1.0 or no correction. 77 Appendix D: Phase Velocities Period (S) Raypaths Land BL SF PR Km/s Km/s Km/s Km/s 16 3347 3.37 3.61 3.55 2.66 18 5213 3.43 3.66 3.54 2.34 20 5611 3.53 3.74 3.89 2.48 22 6306 3.61 3.77 3.86 2.49 25 6739 3.69 3.82 3.89 2.78 29 6275 3.73 3.85 3.93 2.44 33 6440 3.78 3.86 3.90 2.85 40 6377 3.81 3.84 3.93 2.36 45 6395 3.84 3.85 3.93 2.56 50 6257 3.85 3.87 3.94 2.51 55 6163 3.85 3.87 3.93 2.60 59 6276 3.85 3.86 3.91 2.55 67 6189 3.88 3.86 3.91 2.51 78 5856 3.92 3.90 3.90 2.62 Note: BL- Borderland, SF- seafloor, PR – Phase Residual 78 Appendix E: Seismograms Figure E1: Seismograms filtered at 22s are shown for an event which originates from a SE backazimuth (from Chile). The waveforms are notable different as they cross the Borderlands (top) compared to observations at more distant instruments in deeper water (bottom). This complex structure is likely due to scattering and multipathing through the sharp jumps in crustal structure between the continent, Borderland, and deep ocean sea floor. I will investigate with this is causing distortion of the Rayleigh wave across the continental margin. 79