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Volume 13, Number 1
7 December 2012
Q0AN03, doi:10.1029/2012GC004353
ISSN: 1525-2027
Crust and upper mantle structure beneath the Pacific
Northwest from joint inversions of ambient noise
and earthquake data
Lara S. Wagner
Department of Geological Sciences, University of North Carolina at Chapel Hill, CB# 3315, Chapel
Hill, North Carolina 27599, USA ([email protected])
Matthew J. Fouch and David E. James
Department of Terrestrial Magnetism, Carnegie Institution of Washington, Washington, D. C. 20015, USA
Sara Hanson-Hedgecock
Department of Geological Sciences, University of North Carolina at Chapel Hill, CB# 3315, Chapel
Hill, North Carolina 27599, USA
[1] We perform a joint inversion of phase velocities from both earthquake and ambient noise induced
Rayleigh waves to determine shear wave velocity structure in the crust and upper mantle beneath the Pacific
Northwest. We focus particularly on the areas affected by mid-Miocene to present volcanic activity. The joint
inversion, combined with the high density seismic network of the High Lava Plains seismic experiment and
data from the EarthScope Transportable Array, provides outstanding resolution for this area. In Oregon, we
find that the pattern of low velocities in the crust and uppermost mantle varies between the High Lava Plains
physiographic province and the adjacent northwestern Basin and Range. These patterns may be due to the
presence of the Brothers Fault Zone which separates the clockwise rotating northwest Basin and Range from
the relatively undeformed areas further north. Further to the east, the Owyhee Plateau, Snake River Plain
(SRP) and northeastern Basin and Range are characterized by high crustal velocities, though the depth extent
of these fast wave speeds varies by province. Of particular interest is the mid-crustal high velocity sill,
previously only identified within the SRP. We show this anomaly extends significantly further south
into Utah and Nevada. We suggest that one possible explanation is lateral crustal extrusion due to the
emplacement of the high density mafic mid-crustal sill structures within the SRP.
Components: 12,000 words, 10 figures.
Keywords: High Lava Plains; Pacific Northwest; Snake River Plain; ambient noise; surface waves; tomography.
Index Terms: 1744 History of Geophysics: Tectonophysics.
Received 23 July 2012; Revised 19 September 2012; Accepted 15 October 2012; Published 7 December 2012.
Wagner, L. S., M. J. Fouch, D. E. James, and S. Hanson-Hedgecock (2012), Crust and upper mantle structure beneath the
Pacific Northwest from joint inversions of ambient noise and earthquake data, Geochem. Geophys. Geosyst., 13, Q0AN03,
doi:10.1029/2012GC004353.
Theme: Genesis of Continental Intraplate Magmatism: The Example From the Pacific
Northwest, USA
©2012. American Geophysical Union. All Rights Reserved.
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1. Introduction
[2] Since the mid-Miocene, the Pacific Northwest
U.S.A. (PNW) has been subjected to widespread
volcanic activity that is not limited to ongoing
Cascade arc volcanism (where eruptions began in
the region of the present-day arc ca. 45–50 Ma
[Madsen et al., 2006]). Voluminous flood basalt
volcanism began 16.5 Ma with dike eruptions near
Steens Mountain in southeastern Oregon. Following
the Steens eruptions, the locus of volcanism migrated
rapidly northward along a series of N-S rifts, culminating in the Columbia River flood basalts (CRB)
with eruptive centers concentrated near the Idaho/
Washington/Oregon border region [Camp and Ross,
2004] (Figure 1). Over the course of 1.5 Ma, over
220,000 cubic kilometers of basalt were erupted
from the Steens, Chief Joseph, and Monument dike
swarms [Camp and Hanan, 2008]. Silicic volcanic
eruptions began just southwest of the Owyhee
Plateau (OP) 16 Ma [Brueseke et al., 2007]. This
silicic volcanism then migrated away from the OP
in two directions. One track headed northeast to
Yellowstone, roughly parallel to, and at the same
rate as, North American plate motion [e.g., Smith
et al., 2009]. The other progressed northwest to
Newberry Caldera along the Oregon High Lava
Plains (HLP) [e.g., Jordan et al., 2004].
[3] These silicic time progressive volcanic tracks
have been the subject of a great deal of research,
but consensus on their causes remains elusive. This
is partly because of the range of tectonic factors that
exist in the region, some or all of which may have
played a role in their development. These tectonic
factors begin with the potential inheritance of
varied lithospheric structures due to the presence
of terrane boundaries across the area. The location
of the 87Sr/86Sr = 0.706 line (Figure 1) indicates
that Proterozoic North America abuts against
younger provinces along a roughly north-south
trend adjacent to the Oregon-Idaho border, cutting
to the northwest into Washington State and southwest into the Basin and Range [Fleck and Criss,
1985; Ernst, 1988]. Later accreted terranes identified in the area include the Blue Mountain Province
[Dorsey and LaMaskin, 2007; Schwartz et al., 2010]
and the Coast Range Basalt Province [Madsen et al.,
2006], also referred to as the Siletzia terrane [Wells
et al., 1998], which may extend east well into central and eastern Washington and Oregon [Schmidt
et al., 2008; Humphreys, 2009; Schmandt and
Humphreys, 2011; Gao et al., 2011]. In addition,
active ongoing tectonic processes include the
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subduction of the Juan de Fuca and Gorda plates
(remnants of the once much larger Farallon plate)
and the formation and migration of the Mendocino
Triple Junction as the San Andreas transform fault
extended northward. Trench rollback along the
Oregon/Washington coast is progressing at a rate of
35 mm/year relative to a Pacific hot spot reference
frame [Schellart et al., 2008]. Basin and Range
extension, which began in central Nevada 35 Ma,
reached southern Oregon by 22 Ma [Scarberry
et al., 2010]. Basin and Range extension today
ends just south of the High Lava Plains physiographic province as evidenced by the decreasing
amount of displacement along major fault lines in
this area [Trench et al., 2012]. Concurrently, but
independently, beginning at 7 Ma, the transtensional NW trending Brothers Fault Zone developed
across much of the modern High Lava Plains
physiographic province (Figure 1) [Jordan et al.,
2004; Trench et al., 2012]. Geodetic studies
indicate that much of central and eastern Oregon,
Idaho, and northern California and Nevada exhibits
block-like rotation about a pole near the Oregon/
Washington/Idaho border, driven in part by Basin and
Range extension, particularly in southern Oregon
[e.g., McCaffrey et al., 2007]. Trench et al. [2012]
suggest that the Brothers Fault Zone accommodates the greater rates of rotation of blocks south
of the HLP from the comparatively stable blocks
closer to the pole.
[4] Given this tectonic history, a number of
hypotheses have been put forth to explain midMiocene to recent volcanic trends in the Pacific
Northwest. Notably, the similarity between the trend
of the Yellowstone-Snake River Plain (YSRP) and
North American plate motion prompted the notion
that all mid-Miocene to recent volcanism in the
Pacific Northwest can be explained by the impact of
a rising plume head, or the subsequent progression
of the North American plate above a plume tail
[e.g., Morgan, 1972; Pierce and Morgan, 1992;
Geist and Richards, 1993; Hanan et al., 2008;
Pierce and Morgan, 2009; Smith et al., 2009]. A
simple plume model alone, however, cannot explain
the direction and temporal progression of the coeval
HLP volcanic track [e.g., Fouch, 2012]. Some
workers have suggested more complex interactions
between an impacting plume head and other tectonic
factors such as lithospheric topography and/or subduction related corner flow to explain both tracks
[Camp and Ross, 2004; Jordan et al., 2004]. Alternative models for the formation of the HLP volcanic
track that do not require an impacting plume head
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Figure 1. Map of the geologic and tectonic setting of the Pacific Northwest. Brown shaded region indicates the area
covered by the Columbia River and Steens flood basalts (CRSB). Red triangles indicate Holocene volcanoes. Solid white
lines are state boundaries. The dashed red line is the 87Sr/86Sr = 0.706 line [Fleck and Criss, 1985; Ernst, 1988]. Brown
lines indicate boundaries for physiographic provinces from the USGS (http://tapestry.usgs.gov/physiogr/physio.
html), modified to include the boundary of the High Lava Plains (light blue shaded area) from Meigs et al. [2009]
and the Owyhee Plateau from Shoemaker [2004]. Small black dots indicate station locations used in the phase velocity
inversions of Wagner et al. [2010] and Hanson-Hedgecock et al. [2012]. The green arcs delineate the boundaries of the
Brothers Fault Zone (BFZ). Black lines labeled A-A′, B-B′, C-C′, and D-D′ show locations of cross sections shown in
Figures 7–10. The dotted black line in Utah indicates the location of the Wasatch Fault. Other features noted: Idaho
Batholith (IB; light shaded area); Snake River Plain (SRP); Western Snake River Plain (WSRP); western Columbia
Basin (wCB); Owyhee Plateau (OP); Soda Lakes Volcanic Field (SL); Black Rock Desert Volcanic Field (BRD); Great
Salt Lake (GSL); Green River Basin (GRB); Uinta Basin (UB); Newberry Volcano (NB); High Lava Plains (HLP).
have focused instead on processes associated with
deformation of the lithosphere and/or subduction
related tectonics. These include Basin and Range
related extension and rotation [e.g., Cross and
Pilger,
1982],
back-arc
extension
[e.g.,
Christiansen and McKee, 1978], asthenospheric
upwelling due to slab rupture [Liu and Stegman,
2012], and/or rollback of the subducting Juan de
Fuca plate [e.g., Carlson and Hart, 1987; Long
et al., 2012]. While there are aspects of these various models that can be verified only through
imaging to great depths beneath the area, many
discriminators of the underlying processes predicted
by the models can be studied through detailed
imaging of the crust and upper mantle. A number of
recent works have taken advantage of the unprecedented coverage of the EarthScope Transportable
Array (TA), commonly in concert with other coeval
temporary seismic deployment data sets to image
the upper mantle using a variety of methods that
include receiver functions, ambient noise, surface
waves, body waves, and seismic anisotropy [e.g.,
Schutt and Humphreys, 2004; Yuan and Dueker,
2005; Burdick et al., 2008; Roth et al., 2008;
Schutt et al., 2008; Sigloch et al., 2008; Warren
et al., 2008; Yang and Ritzwoller, 2008; Yang
et al., 2008b; Ekström et al., 2009; Long et al.,
2009; Tian et al., 2011; Beghein et al., 2010;
Buehler and Shearer, 2010; Eagar et al., 2010; Lin
et al., 2011; Moschetti et al., 2010a, 2010b;
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Obrebski et al., 2010; Pollitz and Snoke, 2010;
Schmandt and Humphreys, 2010; Schwartz et al.,
2010; Wagner et al., 2010; Xue and Allen, 2010;
Yuan and Romanowicz, 2010; Gao et al., 2011;
James et al., 2011; Obrebski et al., 2011; Schmandt
and Humphreys, 2011; Yang et al., 2011; Yuan
et al., 2011; Hanson-Hedgecock et al., 2012;
Schmandt et al., 2012]. Joint inversions of ambient
noise and earthquake induced Rayleigh wave phase
velocities are particularly well suited to creating
accurate 3-D imagery of the crust and uppermost
mantle, especially when a-priori constraints on
crustal structure are provided (commonly from
receiver functions). Such studies have previously
been performed for the western U.S. [e.g., Stachnik
et al., 2008; Moschetti et al., 2010a; Yang et al.,
2011], but none have focused specifically on the
area of the Pacific Northwest centered on the HLP,
YSRP, and northern Basin and Range. This study is
an attempt to fill that void by performing a joint
ambient noise/earthquake induced Rayleigh wave
phase velocity inversion for 3D shear wave velocity
structure. This work takes advantage not only of the
broad, uniform coverage of the TA, but also the high
density station spacing of the High Lava Plains
seismic deployment, consisting of 118 broadband
stations across southwestern Oregon, Idaho, and
northern Nevada. Our results highlight a number of
structures previously unidentified that can help shed
light on the processes that resulted in the crust and
upper mantle structures within the HLP and YSRP.
2. Data and Methods
[5] For this study, we invert Rayleigh wave phase
velocity dispersion curves to obtain vertical shear
wave velocity profiles at 20,691 different locations in map view. These one-dimensional velocity profiles are subsequently combined to create a
three-dimensional shear wave velocity model. We
define the map view locations at 0.1 intervals
between 38 and 47 N, and 125 and 108 W.
At each location, we compile a dispersion curve
using the phase velocity maps published in Wagner
et al. [2010] (a two-plane wave approach) and
Hanson-Hedgecock et al. [2012] (an ambient noise
approach). Included in these dispersion curves are
the phase velocities at each of 21 periods analyzed
in the aforementioned studies: 8, 10, 12, 15, 20, 25,
28, 30, 33, 35, 40, 45, 50, 58, 66, 77, 91, 100, 111,
125, and 143 s. Below we outline specifics for each
of the steps in this process: 1) determining phase
velocities at each point, 2) creating a 1-D shear wave
velocity reference model, 3) adjusting the reference
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shear wave velocity model at each point to reflect
changes in crustal thickness, and 4) inverting for
shear wave velocity and evaluating regularization
and resolution.
2.1. Determining Phase Velocities at Each
Point
[6] The analysis of earthquake-induced Rayleigh
wave phase velocities of Wagner et al. [2010] was
performed at periods between 25 and 143 s. The
ambient noise inversions for Rayleigh wave phase
velocities of Hanson-Hedgecock et al. [2012] provide constraints on shorter period phase velocities
(with periods ranging between 8 and 40 s). Within
the overlapping range of 25–40 s., Wagner et al.
[2010] analyze four different periods: 25, 28, 33,
and 40 s. Hanson-Hedgecock et al. [2012] calculate
phase velocities for six periods in this range: 25, 28,
30, 33, 35, and 40 s. In Figure 2, we plot the results
of both the ambient noise and earthquake induced
phase velocity inversions for the four overlapping
periods (25, 28, 33, and 40). In general, there is
very good agreement between the results from the
two methodologies, though some differences are
still seen. For this study, we need to assign a phase
velocity for each period to each point in map view.
The question then arises whether to choose one
map over the other, or to average the two results.
In the case of the phase velocities at 25 and 28 s,
the standard deviation of the results reported in
Wagner et al. [2010] are markedly higher than for
the better resolved longer periods. This is largely
because there are fewer data available at shorter
periods for Rayleigh waves induced by earthquakes
due to the more rapid attenuation of higher frequencies over the long distances traveled by the
teleseismic plane waves. Ambient noise induced
Rayleigh waves do not have this limitation. In
comparing the phase velocity maps of the two
methodologies at 25 and 28 s, the primary difference
between the two is in the smoothness of the model:
the earthquake induced Rayleigh wave velocity
maps show smaller scale structures that are not
necessarily reliable and do not appear in the ambient
noise maps. Similar differences are observed at 33 s
but not at 40 s. We decided that at the shorter periods
(25, 28, and 33 s), the ambient noise results were
more conservative than the earthquake Rayleigh
wave results, and therefore use the ambient noise
maps for determining phase velocities at those
periods. At 40 s, we could see no clear reason to
prefer one map over the other, so for this period,
the phase velocities were averaged at each point in
map view.
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Figure 2. Comparison of the results for phase velocities between the ambient noise inversions of Hanson-Hedgecock
et al. [2012] and the earthquake-induced Rayleigh wave inversions of Wagner et al. [2010]. (left) The results of
Wagner et al. [2010]. (right) The results of Hanson-Hedgecock et al. [2012]. Absolute velocities are plotted in
greyscale and are contoured at 0.1 km/sec intervals.
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Figure 3. Average phase velocities for the 2-D phase velocity maps of Wagner et al. [2010] and Hanson-Hedgecock
et al. [2012], and shear wave velocity models with depth. (a) Diamonds indicate RMS average phase velocities used at
each period and error bars are the standard deviations used for our preferred model at their respective periods (see text
for details). Colored circles indicate phase velocities calculated for our preferred reference shear wave velocity model
(shown with heavy black line in Figure 3b). Colored squares show phase velocities calculated for the shear wave
velocity model shown in heavy orange line in Figure 3b. that is the product of a less-damped inversion. (b) Sensitivity
kernels for the periods associated with the same color circles in Figure 3a. In addition to sensitivity kernels, Figure 3b
shows several velocity models discussed in the text. The dashed purple line is the starting model based on the TNA
model from Grand and Helmberger [1984]. The solid purple line is the result of our inversion for shear wave velocity
structure using the aforementioned starting model. The dashed orange line is based on the reference model in Wagner
et al. [2010], and is used as a starting velocity model for the inversions for the reference model for this study. The solid
orange line is the result of a reduced-damping inversion of observed average phase velocities using the dashed-orangeline velocities as a starting model. The solid black line is the result of a heavily damped inversion of the dashedorange-line model using the observed phase velocities, and is our reference starting model.
2.2. Creating a 1-D Shear Wave Velocity
Reference Model
[7] In order to calculate a reference 1-D shear wave
velocity model, we first average the phase velocities
at all points for each period to obtain an average
observed dispersion curve (Figure 3a). As described
above, we used the phase velocities from the ambient noise inversions for periods between 8 and 35 s,
and averaged the results between the ambient noise
results and the earthquake induced surface wave
results for the 40 s phase velocities. For periods
longer than 40 s, we use the earthquake induced
Rayleigh wave phase velocity maps alone. The
starting model for this inversion, modified from
the reference model of Wagner et al. [2010], consists of 6 crustal layers with a total crustal thickness
of 35 km, and 12 upper mantle layers extending to
420 km depth (Figure 3b). Layer thicknesses
increase with depth due to the broadening depth
sensitivity of the longer period surface waves. Ideal
resolution is generally above 150 km depth with
some resolution between 150 and 250 km depth
(see section 2.4 for further discussion). Additional
layers of increasing thickness are included up to
420 km depth in order to accommodate the very broad
sensitivity kernels of the longest period surface waves.
[8] Using this starting velocity model, we invert the
average dispersion curve to determine our reference
1-D shear wave velocity model using the method of
Saito [1988]. The resultant shear wave velocity
model, and the predicted phase velocities for that
model are shown in Figures 3b and 3a respectively. Regularization for this inversion is achieved
by assigning a standard deviation to each phase
velocity observation in the dispersion curve; for
ambient noise phase velocities, we assume a standard deviation of 1% of the phase velocity at a given
period. For periods of 40 s and longer, we use the
average standard deviation from Wagner et al.
[2010] at that period (Figure 3a).
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Figure 4. Moho depth map used to determine crustal
thickness for the starting model in each inversion.
[9] For this study, we seek a reference shear wave
velocity model that fits the observed phase velocities but makes as few assumptions as possible about
existing lithospheric structure, especially with
respect to the presence or absence of a fast mantle
lid. A number of recent studies [e.g., Li et al., 2007;
Abt et al., 2010; Levander and Miller, 2012] have
analyzed variations in the depth to the lithosphereasthenosphere boundary (LAB) across the western
U.S. using both Ps and Sp receiver functions. These
studies each show that the depth to the LAB
(designated as the top of a region of lower shear
wave velocities) varies between 40 and 140 km
within our study area. The inversion described
above does not produce a low velocity zone, but
if we reduce the standard deviation of each of
the phase velocities by an order of magnitude,
the resultant shear wave velocity model (shown in
Figure 3b) does show a subtle low velocity zone
between 70 and 167 km depth. The overlying higher
velocity lid has shear wave velocities below
4.25 km/sec. The phase velocities predicted for this
velocity model are shown in Figure 3a. The fit
of this less-damped model to the observed average
phase velocities is slightly better than our moredamped model, but the differences are small
compared to the average standard deviations at
each period.
[10] In order to test the sensitivity of our results to
the starting model, we repeat this inversion, this
time using starting velocities from the TNA model
of Grand and Helmberger [1984] which include a
high velocity lid (4.4 km/sec) overlying a low
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velocity zone between 50 and 250 km depth. This
inversion produces a low velocity zone starting at
45 km depth, with a maximum lid velocity of
4.29 km/sec. Despite the very different starting
models, the result of this inversion is very similar
to the lower-damping result calculated using the
Wagner et al. [2010] starting model, and is nearly
identical between 70 and 250 km depth. This
suggests that the shear wave velocities across our
study area in the upper most mantle are indeed on
average substantially lower than those found by
Grand and Helmberger [1984]. It also indicates
that the depth to the LAB in our inversion results
(70 km for the lower-damping model using the
Wagner et al. [2010] starting model, 45 km for the
result of the TNA starting model) is dependent
on the starting model, and is therefore not well
resolved.
[11] Given the inconsistency in the geometry of
the low velocity zone, together with the limited
improvement in fit to the observed average phase
velocities for the results that do show a low velocity
zone over the one which does not, we choose to use
the more simple model that makes no assumptions
about the depth to the LAB or low velocity zone.
Also, as discussed in section 2.3, we need to be able
to change crustal thickness at each point in map
view while keeping our starting model as close to
the reference model as possible. The absence of a
mantle lid and low velocity zone greatly facilitates
this process.
2.3. Crustal Thickness Adjustments to the
Reference Shear Wave Velocity Model
[12] The next step is to invert the observed disper-
sion curves at each point in map view for the best fit
shear wave velocity profile. Surface waves are
generally poor at constraining Moho depth so we
adjust the starting model at each point to reflect a
priori constraints on crustal thickness (Figure 4).
The Moho depths shown in Figure 4 were calculated
from a combination of the receiver function results
of Eagar et al. [2011] and from the EARS catalog
for regional Transportable Array stations [Crotwell
and Owens, 2005] (see Wagner et al. [2010] for
details). This is the same Moho map used in
Hanson-Hedgecock et al. [2012], which in turn is a
slightly smoothed version of the Moho map used in
Wagner et al. [2010].
[13] At each point, we first adjust the starting 1-D
model slightly to take into account variations in
crustal thickness. Our adjustment method preserves
the original layer thickness parameterization as much
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Figure 5. RMS average misfit over all periods at each
point for the phase velocities calculated using the final
shear wave velocity model compared to the observed
dispersion curves.
as possible, meaning that the number of layers within
the crust may vary but the number of layers between
the surface and a given depth remains relatively
constant. To accomplish this, we determine which
layer contains the Moho at that point in map view
based on the Moho depths shown in Figure 4.
The part of the Moho-containing layer that lies
above the Moho is added to the overriding layer
thickness, and the part of that layer that is below
the Moho is added to the layer below. The new
thicknesses of the two layers are compared, and, in
order to keep a constant total number of layers
and to avoid overly thick layers, the thicker of the
two is split in half. If layers are moved from crust to
mantle, the new mantle layers are assigned the same
velocities and density as the uppermost mantle layer
in the starting model. If layers are moved from
mantle to crust, the new crustal layers are assigned
the same velocities as the deepest crustal layer in the
starting model.
2.4. Inverting for Shear Wave Velocity:
Evaluating Regularization and Resolution
[14] Using this starting shear wave velocity model
and the phase velocities at that point in map view,
we invert for a 1D shear wave velocity profile at
that point. Regularization of these inversions is
achieved by assigning a standard deviation to each
phase velocity measurement. The goal is to find a
regularization scheme that results in velocity deviations that vary gradually with depth but still allow
significant enough variations to fit the observed
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phase velocities well [e.g., Yang et al., 2008a].
While standard deviations are not available for the
phase velocity maps from Hanson-Hedgecock et al.
[2012], they were provided for the phase velocity
maps of Wagner et al. [2010]. In order to take
advantage of this information, we used the following
regularization scheme for the models shown here:
for periods longer than 40 s, we use a value equal to
half of the standard deviation of the phase velocity
at that point as published in Wagner et al. [2010].
For shorter periods, we take the lesser of 1% of
the phase velocity or the value used for standard
deviation for the 40 s phase velocity. This scheme
was determined by testing a variety of regularization
models, including constant values for standard deviation across all periods, and values weighted solely
by the absolute phase velocities at each period. We
chose this particular regularization method because
it produced velocity deviations that varied gradually
with depth, but still allowed for strong velocity gradients where required to fit the observed phase
velocities. This regularization also takes into account
the varying resolution of phase velocities at different
points in map view.
[15] We evaluate how well our 3D shear wave
velocity model fits the observed phase velocity
maps by calculating predicted phase velocities at
each point, and calculating the RMS misfit at each
point, averaged over all periods. The result of this
test can be seen in Figures 5. For most of the study
area, errors are less than 0.02 km/sec, which compares favorably with the misfits of previous studies
[e.g., Yang et al., 2008b]. We can also assess the
sensitivities of our inversions to depth by looking at
the diagonal of the resolution (R) matrix for each
point in map view. In an ideal inversion, a perfectly
resolved layer with no trade-offs with other layers
would have a diagonal value of 1. Lower values on
the diagonal imply trade-offs with adjacent layers.
We plot the diagonal of the R matrix for each inversion (i.e., each point in map view) in Figure 6a, along
with the mean and standard deviation for each layer.
Layers 1–4 and 9–18 have constant depths and
thicknesses. Layers 5–8 vary somewhat in depth and
thickness due to varying Moho depths. This results in
an increase in the range of diagonal R values,
depending on whether the layer above and below the
Moho ended up being thicker or thinner than the
reference model layer. Thinner layers are more likely
to trade off with adjacent layers than thicker layers,
resulting in a lower value on the diagonal of the R
matrix. Mean R-values are generally above 0.2 for
depths between 5 and 140 km depths, but then
decrease at 140 km depth to between 0.1 and 0.15.
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We observe another decrease at 250 km depth to R
values between 0.01 and 0.07. R-values are also
sensitive to the standard deviations assigned to the
phase velocities. The R-values for layers 4 (15–
20 km), 12 (107–137 km), 13 (137–187 km), and 16
(250–320 km) are shown in Figures 6b–6e. These
show the highest resolution in areas within the High
Lava Plains deployment, and lower resolution at the
periphery where the phase velocities are less well
constrained and have higher standard deviations in
Wagner et al. [2010].
3. Results
[16] Our 3D shear wave velocity model is shown in
Figures 7–10. A number of structures observed in
our results are similar to those seen in previous
tomographic studies of the PNW crust and uppermost mantle. We see the western Columbia Basin,
characterized by very low velocities in the upper
crust underlain by very high velocities (Figure 7)
[Moschetti et al., 2007, 2010a; Gao et al., 2011;
Obrebski et al., 2011; Yang et al., 2011; HansonHedgecock et al., 2012]. We also see the high
velocities of the ancestral Cascades adjacent to the
low velocities of the modern Cascade arc at upperto-mid crustal depths [Moschetti et al., 2007,
2010a; Gao et al., 2011; Obrebski et al., 2011;
Yang et al., 2011; Hanson-Hedgecock et al.,
2012]. Other prominent structures in the crust
include very low velocities associated with the Uinta
and Green River Basins (Figures 7a and 7b) [Yang
et al., 2011], very low velocities along the coast of
Northern California (Figures 7a–7e) [Moschetti
et al., 2007, 2010a; Obrebski et al., 2011; Yang
et al., 2011; Hanson-Hedgecock et al., 2012], and
reduced velocities near the Holocene volcanic centers of Soda Lakes, NV and Black Rock Desert,
UT [Yang et al., 2008b; Moschetti et al., 2010a]
(Figure 7). We see high crustal velocities associated
with the Siletzia terrane to the north (Figures 7a–7e)
that are very similar to those in the study of Gao
et al. [2011]. Gao et al. [2011] likely have better
Figure 6. Resolution matrix diagonal values. (a) Diagonal R-values for all 1-D inversions performed in this
study are plotted with thin gray lines. Red and blue
squares show average R-values for each layer. Horizontal black bars show standard deviations for each layer.
Blue squares indicate layers plotted in Figures 6b–6e.
(b–e) Diagonal R-values in map view for layers 4, 12,
13, and 16 respectively. Depth ranges for each layer
are indicated to the right of each map.
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Figure 7. Shear wave velocity maps at crustal depths and long period gravity anomalies. (a–d) Shear wave velocity
deviations from the starting model at 10, 15, 20, and 25 km depth in color. Absolute velocities are contoured in 0.1 km/
sec increments. Other symbols are the same as in Figure 1. (e) The same, but at 4 km above the Moho depth at each
point. (f) Bouguer gravity anomalies from Kucks [1999] filtered between 100 and 1000 km.
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resolution in this area than this study due to their
incorporation of data from the Wallowa flexible
array deployment in northeastern Oregon and
southeastern Washington that are not included in the
phase velocity inversions of Wagner et al. [2010] or
Hanson-Hedgecock et al. [2012]. In the upper
mantle, we image high velocities associated with the
subducting Juan de Fuca slab to the west, and with
the Wyoming craton to the southeast (Figure 8)
[Moschetti et al., 2007, 2010a; Wagner et al., 2010;
Pollitz and Snoke, 2010; Obrebski et al., 2011; Yang
et al., 2011].
[17] The prominent low velocities in both the crust
and upper mantle associated with the Yellowstone
Caldera at shallow depths, and with the larger
Yellowstone/Snake River Plain at lower crustal and
upper mantle depths are similar to those imaged by
Wagner et al. [2010] and Hanson-Hedgecock et al.
[2012], though some differences exist. Notably,
while Hanson-Hedgecock et al. [2012] found that
very low velocities associated with the Yellowstone
Caldera persisted nearly vertically through the full
thickness of the crust, we image some deviation in
the magnitude and location of the lower crustal low
velocity anomaly (Figure 10). Given that the phase
velocities for periods shorter than 40 s are the same
as those used in the earlier study, the difference
must be due to the addition of phase velocity
information at periods 40 s and longer. The results
of the present study are more consistent with
previous results which also image a transition from
slow to fast crustal material directly below the
Yellowstone Caldera [e.g., Stachnik et al., 2008].
[18] We note that the “bottom” of the low velocity
anomaly in the mantle beneath the YSRP is not well
resolved, and may be due to a decrease in sensitivity
at those depths of the surface wave kernels. As
discussed in section 2.4, while our images provide
details about the upper 250 km of the YSRP system,
they suffer from reduced resolution at depths greater
than about 150 km and therefore cannot provide
insight on the ultimate source depth of the YSRP
volcanic track. Other imaging studies using teleseismic body waves [e.g., Humphreys et al., 2000;
Schutt and Humphreys, 2004; Yuan and Dueker,
2005; Burdick et al., 2008; Roth et al., 2008; Schutt
et al., 2008; Sigloch et al., 2008; Obrebski et al.,
2010; Schmandt and Humphreys, 2010; Xue and
Allen, 2010; James et al., 2011; Schmandt and
Humphreys, 2011; Schmandt et al., 2012] are better
suited to resolve this controversy.
[19] The crust and uppermost mantle throughout
most of southeastern Oregon are dominantly
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characterized by low shear wave velocities [e.g.,
Moschetti et al., 2007; Lin et al., 2009; Moschetti
et al., 2010a; Pollitz and Snoke, 2010; Wagner
et al., 2010; Obrebski et al., 2011; Yang et al.,
2011; Hanson-Hedgecock et al., 2012]. Variations
exist in the locations of the strongest anomalies with
depth. Within the crust (Figures 7c–7e and cross
section B-B′ in Figure 9), the low velocity anomaly
is most prominent in the northernmost Basin and
Range, primarily south of the High Lava Plains
physiographic province. This low velocity anomaly
extends from mid crustal depths to the Moho, and
is most pronounced due north of the California/
Nevada/Oregon border. In this area we observe
shear wave velocity deviations of up to 8%, and
absolute shear wave velocities in the lower crust
below 3.4 km/sec. The area within the High Lava
Plains physiographic province tends to have somewhat less pronounced crustal shear wave anomalies, especially in areas with the most post-CRB
volcanic activity (Figure 9). At upper mantle depths
(Figures 8a–8d), by contrast, the low velocity
anomaly is slightly more pronounced under the
High Lava Plains physiographic province, whereas
the northwestern Basin and Range region exhibits a
less prominent low velocity anomaly. The upper
mantle low velocity anomaly along the main volcanic track of the High Lava Plains appears most
pronounced between 50 and 100 km depth, shallowing from east to west toward Newberry Volcano,
where the HLP low velocity anomalies merge with
low velocity structures associated with arc volcanism (Figure 8 and cross section A-A′ in Figure 9).
[20] In contrast to the low crustal velocities
observed within the High Lava Plains and northwest Basin and Range, we observe a high velocity
anomaly within the crust that extends from the
western margin of the Owyhee Plateau east to the
Wyoming border and Wasatch fault (Figures 7a–
7e). The north-south extent of this anomaly ranges
from the northernmost extent of the western Snake
River Plain south into northern Nevada and Utah.
The depth extent of this anomaly varies, depending
on location. Along the SRP, the anomaly does
not reach the surface and does not extend below
25 km depth (Figure 7e and cross section C-C′ in
Figure 10). Within the Owyhee Plateau, the high
velocity anomaly is particularly pronounced at 10 –
20 km depth (Figures 7a–7c), but also extends down
to the Moho (Figures 7e and 8a and cross section
C-C′ in Figure 10). Areas south of the OP and SRP
also show this high velocity anomaly at various
depths. At 10 – 15 km depth, high velocities are seen
extending into the Lake Bonneville region in
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Figure 8. Shear wave velocity maps at upper mantle depths. Shear wave velocity deviations from the starting model
at (a) 45, (b) 65, (c) 85, (d) 105, (e) 125, and (f) 145 km depth in color. Absolute velocities are contoured in 0.1 km/sec
increments. Other symbols are the same as in Figure 1.
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Figure 9. Zoomed in map and cross sections of the High Lava Plains region. The background velocity anomalies for
the map-view plot are at 25 km depth. Upright triangles are rhyolitic volcanoes color-coded by age from Meigs et al.
[2009]. Inverted pink triangles indicate locations of Holocene volcanism. Cross-sections show shear wave velocity
deviations from the starting model. Heavy black line indicates the Moho used in this inversion, and the thin black line
shows overlying topography. The bottom axis of each of the cross sections shows distance in km along the transect,
whereas the numbers above the topography indicated the longitude.
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Figure 10
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northwestern Utah. Below 20 km depth, this anomaly
broadens across northern Nevada, although the highest amplitude anomalies remain in northeasternmost
Nevada and northwestern Utah. The area in northwestern Utah and northeastern Nevada with the
thickest and most prominent high velocity anomaly
corresponds spatially with a regional long-wavelength
Bouguer gravity high (Figure 7f) [Kucks, 1999;
DeNosaquo et al., 2009]. The D′-D cross-section in
Figure 10 shows a profile across the gravity anomaly
map from Figure 7f. This cross-section shows a
striking spatial correlation between the topographic
lows of the SRP and Great Salt Lake Basin area with
relative Bouguer gravity highs. In between the two is
an area of higher topography with somewhat lower
gravity anomalies. In the seismic cross-section, this
region represents the transition from the shallower,
thinner high velocity anomaly within the SRP to the
thicker, deeper high velocity anomaly further south.
It is also underlain by the strong mantle low velocity
anomaly that extends the length of the YSRP.
[21] In this paper, we will focus on two major
features: the low velocity structures associated with
the High Lava Plains (Figures 7–9), and the high
velocity crustal anomaly that we observe beneath
the SRP, Owyhee Plateau, and parts of northern
Utah and Nevada (Figures 7, 8, and 10). Both of
these observations are roughly consistent with most
previously published velocity models, but the
details and tectonic implications of each have not
previously been fully investigated.
4. Discussion
4.1. The High Lava Plains Low Velocity
Anomaly
[22] Two aspects of the low velocity anomaly across
southeastern Oregon are particularly striking: (1) the
most pronounced low velocities in the lower crust
lie south of the main trend of HLP volcanism; and
(2) the most pronounced low velocities in the
uppermost mantle are not located below the crustal
low velocities, but rather to the north beneath the
HLP physiographic province. These results are
not entirely unexpected. The uppermost mantle
anomalies observed are similar to those in Wagner
et al. [2010]. The very low shear wave velocities
in the crust immediately above the Moho in the
northwestern Basin and Range are consistent with
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the results of Hanson-Hedgecock et al. [2012].
These results are also consistent with receiver
functions from this area that have mid-crustal negative polarity arrivals consistent with a decrease in
velocity with depth [Eagar et al., 2011]. Receiver
function H-k stacking indicates unusually high
Poisson’s ratio in this area as well [Eagar et al.,
2011] which are also consistent with our observed
low shear wave velocities.
[23] Combined with observed high heat flow
[Blackwell and Richards, 2004] and low resistivity
[Patro and Egbert, 2008] throughout the area, one
possible explanation for the very low velocities in
the lower crust of the northwestern Basin and Range
is the presence of small degrees of partial melting.
For crustal compositions, an 8% reduction in shear
wave velocities would correspond to a melt fraction
of 3% [Watanabe, 1993; Takei, 2000; Caldwell
et al., 2009], assuming the entire velocity anomaly
is due to partial melt alone. This would suggest the
presence of up to 3% partial melting in the lower
crust beneath the northwestern Basin and Range,
which is roughly consistent with the 2% found by
Eagar et al. [2011] based on the high Vp/Vs ratio
determined using receiver function H-k stacking.
If this anomaly is indeed due to partial melting,
the location of this lower crustal partial melting
is notably not centered on the area in the High
Lava Plains with the most significant amount of
post-flood-basalt volcanism. Eagar et al. [2011]
suggest this may be evidence that these represent
zones of “undrained” partial melt in the crust,
whereas the area further north along the HLP physiographic province had erupted most or all of the
partial melt previously present in the crust.
[24] One possible explanation for why the melting
was able to erupt further to the north but not within
the northern Basin and Range is the Brothers Fault
Zone (BFZ), which may have acted as a conduit
for the magma to escape. The BFZ began to develop
sometime between 7.5 and 5.7 Ma, synchronously
with, but separately from Basin and Range extension further south [Jordan et al., 2004; Trench et al.,
2012]. This roughly coincides with a change in
relative plate motions as suggested by Jordan et al.
[2004], resulting in a block rotation of much of the
study area about a pole located near the Oregon/
Washington/Idaho border [e.g., McCaffrey et al.,
2007]. The NW trending faults of the BFZ delineate
Figure 10. Perspective plot of study area with cross-sections showing the YSRP and the crustal high velocity anomaly extending into Nevada and Utah. Colors and symbols on cross-sections the same as in Figure 9. Bottom cross
section shows profile through the filtered gravity anomaly from Figure 7f in purple superimposed on the topography.
Other symbols are the same as in Figure 9.
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a small circle about this pole, separating Basin and
Range extension from the relatively undeformed
block closer to the pole. The spatial correlations
between the BFZ, the main volcanic trend of the
HLP, and the upper mantle low velocity zone,
together with the lack of correlation with the most
profound lower crustal low velocity anomalies
suggest two possible hypotheses: (1) The BFZ
formed in this location because of a pre-existing
zone of weakness, seen today as a low velocity zone
in the uppermost mantle, resulting in a focusing of
volcanism along the HLP; or (2) the formation of the
BFZ along the northern margin of the Basin and
Range resulted in an area of shearing, weakening,
and/or heating in the uppermost mantle, as indicated
by the low velocities and apparent lack of mantle
lithosphere observed today. The pattern of volcanism along the HLP indicates four major pulses of
volcanic activity between 7.8 and 2 Ma [Trench
et al., 2012]. These volcanic events may have produced a temporary localized weakening of the crust
which allowed the BFZ to develop in that area.
However, after melt extraction, the residual more
mafic crust would result in an increase in crustal
strength [Brace and Kohlstedt, 1980; Trench et al.,
2012], perhaps contributing to the westward
migration of silicic volcanism along the HLP. Long
et al. [2012] suggest that the temporal pattern of the
HLP silicic volcanism can be explained by patterns
of trench roll-back and flow around the southern
edge of the subducted slab. They argue that trench
roll-back, together with a steepening of the subducted slab, would have initiated 20 Ma, resulting
in a pulse of upwelling that produced the midMiocene flood basalts and thinning of the mantle
lithosphere across much of the area. Ongoing trench
retreat, together with toroidal flow around the southern edge of the subducted slab produced a W-to-E
flow field that helps to explain both the migration of
HLP volcanism and the presence of large magnitude
E-W oriented SKS splitting [Long et al., 2009].
While these processes also produced small degrees of
partial melting in the lower crust in areas south of
the HLP, the quantities of magma were sufficiently
small, or conduits to the surface were lacking, such
that magma extraction did not occur, and the lower
crust in this region remains today as a profoundly low
velocity zone in the northwestern Basin and Range.
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in crustal velocity maps of a number of previous
tomographic studies [e.g., Pollitz and Snoke, 2010;
Yang et al., 2011]. The notable exception is
Moschetti et al. [2010a, 2010b], though other more
recent works by co-authors do show this feature
[e.g., Yang et al., 2011]. The principle difference
between Moschetti et al. [2010a, 2010b] and other
studies is the inclusion of Love waves and radial
anisotropy. This might suggest that our observed
fast shear wave velocities in the upper mantle are due
to radial anisotropy, but the results of Moschetti et al.
[2010b] do not show a corresponding high radial
anisotropy anomaly in their upper mantle models in
this area. Evidence for the presence of high velocity
material in the crust across the area can also be
seen in other seismic observables. The high velocity
anomaly directly above the Moho (Figure 7e)
near the OP is reminiscent of Eagar et al. [2011,
Figure 10], which shows the amplitudes of Ps
converted phases from Moho depths: the areas in
which we observe high supra-Moho velocities
exhibit particularly low Ps converted phase amplitudes in the Eagar study, consistent with a decreased
seismic impedance contrast across the Moho due to
elevated lower crustal velocities, as well as perhaps
a sharply dipping Moho.
[26] The Owyhee Plateau (OP) itself is seen in our
models as having high velocities from the surface to
the Moho. A sharp boundary between the OP and
areas to the west can be seen in Moho thicknesses,
as the OP has a distinctly thicker crust than the
adjacent HLP [Eagar et al., 2011]. Unlike the surrounding areas, the OP experienced comparatively
little volcanism since the mid Miocene, with most
of the volumetrically significant volcanism occurring
along the western and northern margins [Shoemaker,
2004]. Hanson-Hedgecock et al. [2012] suggest that
this anomaly could be due to Precambrian lithosphere that was further depleted during the eruption
of the Columbia River flood basalts. We note that
unlike adjacent areas (discussed below), the OP is
characterized by comparatively high elevations. This
is consistent with the increased crustal thickness in
the area, but may limit the amount of dense mafic
restite that could present in the mid and lower
crust without isostatically reducing the surface
topography.
[27] The portion of the high velocity anomaly that
4.2. The Crustal High Velocities of the
Snake River Plain/Owyhee Plateau/Great
Salt Lake
[25] The broad extent of the high velocity crustal
anomaly in our model is similar to anomalies seen
has been most investigated is the mid-crustal feature
within the bounds of the Snake River Plain. In our
models, this anomaly is clearly not just limited to the
SRP, but extends to the south into northwestern
Utah where it thickens substantially (Figure 9, cross
section D-D′). Previous studies have focused solely
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on the portion of this anomaly located within the
SRP by investigating cross-sections that either
entirely or mostly exclude the region of the high
velocity anomaly we observe south of the SRP.
For example, Peng and Humphreys [1998] found
evidence for a high velocity mid-crustal layer that
does not extend beyond the margins of the SRP.
The seismic array used, however, was deployed in
a transect north and east of the Wasatch fault,
beyond the easternmost extent of our observed high
velocities. Their work builds on the earlier reflection/
refraction study along a similar transect by Sparlin
et al. [1982] who found evidence for a high velocity, high density mid-crustal layer localized within
the margins of the SRP. Similar results have been
found in subsequent seismic studies [e.g., Schutt
et al., 2008; Stachnik et al., 2008], but these also
only investigate the portion of the Snake River Plain
northeast of the Wasatch Fault.
[28] The most common explanation for the mid-
crustal SRP high velocity anomaly is a gabbroic
sill or series of sills [Shervais et al., 2006] left
behind by the voluminous silicic eruptions along the
SRP [Sparlin et al., 1982; McQuarrie and Rodgers,
1998; Peng and Humphreys, 1998; Schutt et al.,
2008; Stachnik et al., 2008; DeNosaquo et al., 2009;
McCurry and Rodgers, 2009; Rodgers and McCurry,
2009]. This high density structure is responsible for
the isostatic subsidence of the SRP relative to the
surrounding Basin and Range [McQuarrie and
Rodgers, 1998]. However, a number of studies have
also suggested that some amount of crustal material
must have been extruded out of the SRP into adjacent
areas. McQuarrie and Rodgers [1998] modeled the
structure of the crust in this region by looking
at isostatically induced flexure along the northern
margin of the SRP. Their results indicate that isostatic
compensation of a mid crustal high density structure
occurs in the lower crust rather than the upper mantle,
accommodated by lateral extrusion of crustal material to areas north and south of the SRP. A similar
conclusion is reached by McCurry and Rodgers
[2009] and Rodgers and McCurry [2009] who use
petrological and kinematic constraints respectively to
argue for the extrusion of restite and densified crust
into areas adjacent to the SRP. Yuan et al. [2010]
argue that crustal material has been extruded, in
their case to the north, eminating from the northeastern end of the SRP near the Heise caldera field.
Their conclusions are based on receiver function
derived crustal thicknesses that indicate increased
Moho depths, not just along the SRP volcanic track,
but also to the north in the area described above.
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DeNosaquo et al. [2009] investigated the mid-crustal
sill structure by combining gravity data with constraints from seismic imaging and seismicity patterns
among others. They focus on the same cross section
as Peng and Humphreys [1998] that does not intersect the southern extension of our high velocity
anomaly but does however include the northeasternmost corner of the gravity high. Based on the extent
of their observed gravity anomaly, they argue for a
slight widening of the proposed sill structure caused
by lateral extrusion of the mid-crustal sill accommodated by normal faults present in the northern
Basin and Range. They also attribute a southward
deflection of the seismogenic parabola surrounding
the YSRP to this widened sill, arguing that the seismicity is the product of flexure induced by sill-related
isostatic subsidence.
[29] We propose that a lateral extrusion of high
density, high seismic velocity material from the
central Snake River Plain may be an explanation for
our observations in northwestern Utah and northeastern Nevada. The adjacent areas of the SRP,
including the Bruneau-Jarbidge and Twin Falls
volcanic centers, produced several times the volume
of silicic magmatism observed at Yellowstone
[Perkins and Nash, 2002; Bonnichsen et al., 2008],
and presumably also produced correspondingly
more gabbroic residual. No difference in crustal
thickness or in the thickness of the mid-crustal high
velocity zone within the SRP is seen, however,
consistent with mass removal due to lower crustal
flow. The area of the northeastern Basin and Range
due west of the Wasatch fault is known to have
undergone a great deal of extensional thinning,
beginning in the late Cretaceous [Wells and Hoisch,
2008] and continuing into the Cenozoic [Egger
et al., 2003] (Figure 4). This extension could have
helped accommodate the influx of additional material, similar to, but on a larger scale than what was
proposed by DeNosaquo et al. [2009]. This influx of
high density material could account for the relatively low topography across the area, and may
explain the relative quiescence of the seismic
parabola surrounding the SRP west of the Wasatch
fault [Anders et al., 1989]. While thinner crust might
be used to explain lower topography as well, we
note that similarly shallow or shallower Moho
depths are seen across much of the Basin and Range
(Figure 4) but the area of low topography is
restricted to the area with the increased Bouguer
gravity anomaly. Future modeling of the observed
gravity and seismic anomalies is required to
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determine the density, volume, composition, and
origin of this high velocity crustal material.
5. Summary
[30] Our inversions for the detailed shear wave
velocity structure of the crust and upper mantle in
the Pacific Northwest show low velocities in the
crust and upper mantle in the High Lava Plains, and
high velocities in the crust across the Owyhee
Plateau, Snake River Plain, and northeastern Basin
and Range. The anomalies in southeastern Oregon
indicate that the lower crust in the northwestern
Basin and Range may contain up to 3% partial melt,
whereas the adjacent High Lava Plains physiographic province that has experienced more volcanism since the mid-Miocene shows no such
anomaly. This may be due to the presence of the
Brothers fault zone, which allowed partial melts to
reach the surface in this area.
[31] The crustal high velocity anomalies observed
within the Owyhee Plateau, Snake River Plain, and
northeastern Basin and Range vary in depth and
intensity. The Snake River Plain high velocities are
constrained to the mid-crust, whereas the Owyhee
Plateau and areas of northwestern Utah have high
velocities that extend up to, and perhaps even into,
the mantle. The high velocity crustal anomaly in the
northeastern Basin and Range, particularly in
northwestern Utah, is well correlated with a
regional Bouguer gravity high. This high velocity
crustal anomaly may be due to lateral flow of mafic
restites from the Snake River Plain into this very
extended region.
Acknowledgments
[32] Data for this project come in part from the High Lava Plains
seismic experiment, funded by NSF Continental Dynamics EAR0507248 (MJF) and EAR-0506914 (DEJ), as well as from EarthScope’s USArray Transportable Array (http://earthscope.org).
LSW was supported in part through NSF grant EAR-0809192.
We would like to thank the wonderful people at the IRIS PASSCAL Instrument Center and at the IRIS Data Management Center
for their ongoing support. We would also like to thank Jenda
Johnson, Steven Golden, and the many people who helped install
and service the 118+ stations in this deployment. Also thanks to
Allen Glazner and John Bartley for their helpful suggestions on
Basin and Range tectonics. Finally, we wish to thank the many
ranchers and other landowners who generously and freely agreed
to host seismic stations for the HLP experiment on their property.
We would also like to thank Brandon Schmandt and one anonymous reviewer for their helpful comments and suggestions.
10.1029/2012GC004353
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