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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
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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
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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
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List of Figures
Figure 1.1 Simplified map of the Borderlands
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Figure 1.2 Crust and Lithospheric variation across my study area
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Figure 1.3 Seafloor magnetic isochron map
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Figure 2.1 ALBACORE OBS deployment area
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Figure 2.2 Behavior of amplitudes with increasing depths
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Figure 2.3 S-Wave velocity variation with depth
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Figure 2.4 Phase Velocity variation with period (Starting model)
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Figure 2.5 Behavior of amplitudes after applying amplitude correction
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Figure 2.6 Azimuthal distribution of earthquakes
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Figure 2.7 Ray path coverage for the study area
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Figure 3.1 Grid node configuration
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Figure 4.1 Average phase velocity as a function of period in three regions
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Figure 4.2 Average phase velocity as a function of period in five regions
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Figure 4.3 Average phase velocity as a function of period in seven regions
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Figure 4.4 Anisotropy deep seafloor
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Figure 4.5 Anisotropy mid age seafloor
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Figure 4.6 Anisotropy outer Borderland
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Figure 4.7 Anisotropy inner Borderland
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Figure 4.8 Anisotropy North Garlock region
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Figure 4.9 Anisotropy South Garlock region
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Figure 4.10 Anisotropy West San Andreas region
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List of Tables
Table 2.1 P-wave, S-wave, and density variation in study region
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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
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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.
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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
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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
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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.
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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).
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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.
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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
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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
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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
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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
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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),
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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
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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
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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-
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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
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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). At periods above 55s, fast wave direction of the anisotropy
at deep ocean, mid ocean and outer borderland is trending NW – SE direction (aligned
with Pacific plate motion) exhibits transition from oceanic lithosphere to asthenosphere
and also mark the lithosphere-asthenosphere boundary. The Moho depth beneath the Los
Angeles Basin is 30 km. The lithospheric thickness beneath the inner Borderland is
approximately 50 km.
55
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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