Download Seismic velocity structure and anisotropy of the Alaska subduction

Journal of Geophysical Research: Solid Earth
RESEARCH ARTICLE
10.1002/2014JB011438
Key Points:
• Surface wave tomography reveals
variations within the Alaska
subduction zone
• Maps of azimuthal anisotropy show
the influence of the Yakutat block
• Anisotropy results can be
reconciled with some, but not all,
SKS observations
Supporting Information:
• Text S1–S3 and Figures S1–S15
• Readme
Correspondence to:
C. Tape,
[email protected]
Citation:
Wang, Y., and C. Tape (2014), Seismic
velocity structure and anisotropy
of the Alaska subduction zone
based on surface wave tomography,
J. Geophys. Res. Solid Earth, 119,
8845–8865, doi:10.1002/2014JB011438.
Received 7 JUL 2014
Accepted 8 OCT 2014
Accepted article online 18 OCT 2014
Published online 2 DEC 2014
Seismic velocity structure and anisotropy of the Alaska
subduction zone based on surface wave tomography
Yun Wang1 and Carl Tape1,2
1 Geophysical Institute, University of Alaska, Fairbanks, Alaska, USA, 2 Department of Geology and Geophysics, University
of Alaska, Fairbanks, Alaska, USA
Abstract
Southcentral Alaska is a complex tectonic region that transitions from subduction of Pacific
crust to flat slab subduction—and collision—of overthickened Yakutat crust. Because much of the Yakutat
crust has been subducted, seismic imaging is needed in order to understand the crustal and upper
mantle structural framework for this active tectonic setting. Here we use teleseismic Rayleigh waves to
image large-scale variations in shear wave structure. Our imaging technique employs a two-plane wave
representation with finite frequency sensitivity kernels. Our 3-D isotropic model reveals several features: the
subducting Pacific/Yakutat slab, slow wave speeds characterizing the onshore Yakutat collision zone, slow
wave speeds of the Wrangell subduction zone, and a deep tomographic contrast at the eastern edge of the
Pacific/Yakutat slab. We produce anisotropic phase velocity maps that exhibit variations in the fast
direction of azimuthal anisotropy. These maps show the dominance of the Yakutat slab on the observed
pattern of anisotropy. West of the Yakutat slab the fast directions are approximately aligned with the plate
convergence direction. In the region of the Yakutat slab the pattern is more complicated. Along the
margins of the slab the fast directions are roughly parallel to the margins. We identify notable differences
and similarities with published SKS splitting measurements. Integrative modeling using 3-D anisotropy
models and different seismic measurements will be needed in order to establish a detailed 3-D anisotropic
velocity model for Alaska. This study provides a large-scale starting point for such an effort.
1. Introduction
Seismic waves are sensitive to three-dimensional (3-D) variations in Earth’s structure. From a seismological
perspective, Earth structure is characterized in terms of density and the elastic tensor. Measurements of seismic waves can be used within formal inverse problems to infer the unknown structure of Earth; this imaging
procedure is known as seismic tomography. Surface waves (Love and Rayleigh) are strongly sensitive to variations in shear wave speed and have long been used to infer isotropic and anisotropic variations in global
Earth structure [e.g., Woodhouse and Dziewonski, 1984; Ekström and Dziewonski, 1998; Romanowicz, 2003;
Trampert and Woodhouse, 2003; Forsyth and Li, 2005].
In this study we use frequency-dependent measurements of Rayleigh waves between periods of 22 and
143 s in order to image large-scale variations of shear wave speeds within the Alaska subduction zone.
Our analysis incorporates four previous temporary PASSCAL (Program for Array Seismic Studies of the
Continental Lithosphere) seismic deployments, as well as all existing broadband station data in Alaska.
Our tomographic images provide the first surface-wave-derived images of the Alaska subduction zone,
and they complement several studies using local or teleseismic body waves.
Anisotropy in the Alaska subduction zone has been studied from the perspectives of (1) SKS shear-wave
splitting [Christensen and Abers, 2010; Hanna and Long, 2012], (2) local earthquakes [Tian and Zhao, 2012],
(3) numerical models of mantle flow using realistic slab geometries [Jadamec and Billen, 2010], and (4) fabric
analysis of exhumed upper mantle rocks [Mehl et al., 2003]. Our anisotropic phase velocity maps provide a
critical piece within the discussion of anisotropic structure and mantle flow within the Alaska subduction
zone (section 5).
1.1. Tectonic Setting
Subduction has occurred in southern Alaska since at least 160 Ma and perhaps as early as 215 Ma, based
on dated arc rocks, stratigraphic units, and plate reconstructions [Fisher and Magoon, 1978; Jarrard, 1986;
Rioux et al., 2010]. Subduction was interrupted by the accretion of the ancient Talkeetna arc some time
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©2014. American Geophysical Union. All Rights Reserved.
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Figure 1. Reference map for terms discussed in the text. Triangles denote active volcanoes (www.avo.alaska.edu).
Labels: F = Fairbanks, A = Anchorage, CI = Cook Inlet, KP = Kenai peninsula, WV = Wrangell volcanoes, and TF = Talkeetna
formation. Blue lines mark the lateral extent of deep slab seismicity, which we digitized from the Alaska Earthquake
Center catalog (M ≥ 1, 1990–2013); note the segment beneath the Wrangell volcanoes (WV), which was digitized from
Figure S4.
between 160 Ma and 100 Ma [Burns, 1985; Wallace et al., 1989; Plafker et al., 1989]. The Talkeetna formation (e.g., Figure 1) and associated ultramafic rocks comprise this Jurassic arc, which is one of only two arcs
on Earth where the ancient Moho, separating upper mantle rocks from lower crustal rocks, is visible at the
surface [Burns, 1985; Mehl et al., 2003]. Furthermore, the subsurface expression of the arc is manifest as the
Southern Alaska Magnetic High [Saltus et al., 1999, 2007].
One of the largest perturbations to subduction is one that is ongoing today: the collision of the Yakutat
block in southern Alaska at the eastern margin of the Aleutian-Alaska subduction zone [Plafker et al., 1978].
This collisional process is inferred to have initiated at approximately 24 Ma [Haeussler et al., 2008; Benowitz
et al., 2011]. This wedge of overthickened crust, presumably an oceanic plateau [Christeson et al., 2010],
is considered to be responsible for the uplift of the Alaska Range, the rapid uplift and remarkable relief
of the coastal Saint Elias mountains, the propagation of intraplate deformation several hundred kilometers from the trench, the broad flattening of the subducting slab, the formation of the Northern Foothills
fold-and-thrust belt, the extinguishing of the eastern Aleutian arc, the activity of the Wrangell volcanic
arc, the lateral flow at the edge of the Pacific/Yakutat slab, and rupturing of the oceanward Pacific crust
(Mw ≤ 7.8 events during 1987–1992) [Plafker et al., 1978; Richter et al., 1990; Hwang and Kanamori, 1992;
Eberhart-Phillips et al., 2006; Bemis and Wallace, 2007; Ridgway et al., 2002, 2007; Fuis et al., 2008; Haeussler,
2008; Jadamec and Billen, 2010; Christensen and Abers, 2010].
Figure 2 identifies the surface and subsurface outline of the Yakutat block [Eberhart-Phillips et al., 2006].
(The “Yakutat block” [Plafker et al., 1978; Fletcher and Freymueller, 1999; Gulick et al., 2007] has alternatively
been referred to as “Yakutat terrane” [Plafker et al., 1989; Worthington et al., 2012; Christeson et al., 2010] and
“Yakutat microplate” [McCalpin et al., 2011; Chapmand et al., 2011; Benowitz et al., 2011]. The term “Yakutat
slab” [Eberhart-Phillips et al., 2006; Benowitz et al., 2011; Finzel et al., 2011], which we also use, emphasizes the
subsurface portion of the Yakutat block that was imaged by Ferris et al. [2003] and Rondenay et al. [2008].)
The Yakutat block is exposed at the surface, but only offshore; its onshore portion is underthrusting the
North America plate and is inferred from seismic imaging, seismicity, and tectonic models. We use the labels
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Figure 2. Active tectonic setting of the Aleutian-Alaskan subduction zone, Southcentral Alaska. See caption of Figure 1
for description of some features. Red lines denote active faults [Koehler et al., 2012]. Circles: A = Anchorage, MM =
Mount McKinley, F = Fairbanks, and J = Juneau. Colored image SW of Anchorage is the basement surface of Cook Inlet
basin, with a maximal depth of 7.6 km [Shellenbaum et al., 2010]. Yellow outline denotes the (minimum) subsurface
extent of the Yakutat slab [Eberhart-Phillips et al., 2006]. The five edges Y1, …, Y5 are used for reference in the text
(section 1.1). Black outline marks the aftershock zone of the 1964 Mw 9.2 earthquake. Plate labels: NA = North America,
PA = Pacific, YK = Yakutat block; arrows indicate PA motion relative to NA from the model of Bird [2003].
Y1, …, Y5 throughout this paper to refer to the boundaries of the block. At its southern extent (north of Y1),
the Yakutat block is >30 km crust and includes several kilometers of sedimentary strata [Christeson et
al., 2010; Worthington et al., 2010, 2012]. Based on marine seismic surveys, Brocher et al. [1994] identified
Yakutat crust in the subsurface of the northern Gulf of Alaska. At its northern extent (near Y4), the Yakutat
block is imaged as thick subducting crust beneath the Alaska Range [Ferris et al., 2003; Rondenay et al., 2008].
The boundary Y3 marks a sharp transition between slab seismicity in the west to none in the east. This transition is a key imaging target of our study. The western boundary (Y5) is not well defined and could be a
gradual transition [Eberhart-Phillips et al., 2006]; however, recent imaging results show a sharp transition
from Yakutat to Pacific crust at a point southeast of Anchorage [Kim et al., 2014].
The Wrangell volcanoes are situated north of the Y2 edge of the (subsurface) Yakutat block (Figure 2). Relatively limited deep seismicity is sufficient to reveal a Benioff zone of the Wrangell subduction zone down
to a depth of 120 km [Stephens et al., 1984; Page et al., 1989; Fuis et al., 2008]. It is unclear whether this
subducting crust is as thick as the Yakutat block.
West of the inferred Yakutat boundary is the Cook Inlet fore arc basin, which records a marine depositional
history of 180 Ma in 16 km of strata [Fisher and Magoon, 1978]. Toward the west of the Yakutat block the
Aleutian-Alaska subduction zone exhibits shorter trench arc distances and active volcanism, marking a
departure from the “perturbed” subduction by the Yakutat block in the east.
1.2. Previous Work
Previous seismic studies of the Alaska subduction zone have used P and S body wave travel times, either
from local or global events [Kissling and Lahr, 1991; Zhao et al., 1995; Eberhart-Phillips et al., 2006; Qi
et al., 2007a, 2007b; Tian and Zhao, 2012]. The most comprehensive data set used thus far is that of
Eberhart-Phillips et al. [2006], which incorporated stations from the BEAAR (Broadband Experiment Across
the Alaska Range) experiment (Figure 3) as well as active source data such as from the EDGE and TACT
(Trans-Alaska Crustal Transect) experiments [Wolf et al., 1991; Fuis et al., 1991; Brocher et al., 1994; Fuis et al.,
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Figure 3. Broadband seismic stations used in this study, described in section 2.1. Permanent stations (as of 2012) =
white, ARCTIC = red, BEAAR = cyan, MOOS = magenta, STEEP = green. Open black triangles show the broadband stations
that were excluded due to poor data quality or suspected instrument response errors or because they were outside the
study region. The road system in Alaska is plotted as black thin lines and is mostly covered by (temporary) stations from
ARCTIC, BEAAR, and MOOS.
2008]. They used travel times from local earthquakes—both from within the crust and within the slab—to
jointly invert for hypocenters and 3-D structure variations in VP and VP ∕VS . Their model revealed the subducting slab, as well as variations within the crust, such as the crustal thickness and also the presence of
large sedimentary basins.
Previous studies of anisotropy in Alaska have been primarily from the analysis of SKS shear wave splitting
[Christensen and Abers, 2010; Hanna and Long, 2012; Song and Kawakatsu, 2013; Perttu et al., 2014]. These
studies have shown a striking transition in the observed fast-splitting direction at the northern extent of
the subducted Yakutat block. However, one challenge with SKS splitting measurements is in interpreting
the volumetric region to which the measurement is sensitive [e.g., Sieminski et al., 2008], which is critical for
interpreting whether the splitting occurs within the subslab mantle, the slab oceanic lithosphere, the mantle
wedge, or even the crust.
Tian and Zhao [2012] performed a tomographic inversion that estimated P wave anisotropy using P and
S travel times from local earthquakes [Wang and Zhao, 2008]. They experimented using starting models
with and without a slab derived from seismicity and tomography [Zhao et al., 1995]. Their emphasis was in
imaging the detailed VP and VS isotropic structure, as well as the P wave anisotropic structure. The P wave
anisotropy revealed a depth-dependent transition between trench-normal and trench-parallel fast directions across the subducting Pacific slab. At 65 km depth the fast directions of P waves is quite similar to
the fast directions from SKS splitting results: trench normal within the subducting Pacific plate and trench
parallel in the mantle wedge beneath the North American plate.
Previous studies have used body wave travel times of local earthquakes [Eberhart-Phillips et al., 2006; Tian
and Zhao, 2012]. In our study, we image the Alaska subduction zone using frequency-dependent surface
waves from earthquakes far outside Alaska. Our technique allows us to image a much larger domain than
previous studies, both in terms of depth extent and lateral extent. This is particularly important for trying to
understand the pattern of anisotropy within and below the subducting slab, which are regions with poor
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to no sensitivity from local earthquakes. Our resolvable length scales, however, are longer than those in
Eberhart-Phillips et al. [2006] and Tian and Zhao [2012], such that we can only interpret the large-scale
features of the region.
We present a VS model for the Alaska subduction zone that provides insights into the structure and dynamics at the edge of a subduction and collisional setting. The most striking features in our tomographic
images include (1) the subducting Pacific-Yakutat slab, (2) the shallow low-wave-speed region in southeastern Alaska associated with Yakutat collision, and (3) the influence of the subsurface Yakutat block on
seismic anisotropy.
2. Seismic Waveforms for the Tomographic Inversion
Our tomographic inversion uses vertical component recordings of Rayleigh waves from earthquakes far
from Alaska, between epicentral distances of 30◦ and 120◦ . Next we describe our selection of earthquakes
and waveforms for the tomographic inversion.
2.1. Station Coverage and Earthquake Selection
Our study uses all available permanent broadband stations in Alaska (as of 2012), as well as stations from
four temporary PASSCAL experiments: BEAAR (Broadband Experiment Across the Alaska Range, active during 1999–2001) [Meyers et al., 1999; Ferris et al., 2003], ARCTIC (Alaska Receiving Cross Transects for the Inner
Core, active during 2005–2007) [Song et al., 2005; Lindner et al., 2010], MOOS (Multidisciplinary Observations Of Subduction, active during 2007–2009) [Christensen et al., 2008; Li et al., 2013], and STEEP (SainT Elias
Erosion and tectonics Project, active 2005–2010) [Hansen et al., 2005; Christeson et al., 2013]. We use waveforms from 168 stations (Figure 3), more than half of which were from temporary deployments. We note
that previous tomographic studies in Alaska [e.g., Zhao et al., 1995; Eberhart-Phillips et al., 2006] have relied
primarily on travel time data, not waveform measurements, and therefore these studies were able to use
short-period stations in addition to broadband stations. But no previous studies have used MOOS, ARCTIC,
or STEEP stations within an Alaska-scale inversion, as far as we know.
We selected earthquake sources from the Global Centroid-Moment-Tensor catalog [Dziewonski et al., 1981;
Ekström et al., 2012] with the following criteria: origin times between 2000 and 2012, magnitudes greater
than 6.0, depths less than 100 km, and epicentral distances 30◦ ≤ Δ ≤ 120◦ from the center of the region,
taken to be (−147◦ , 62◦ ). The minimum epicentral distance of 30◦ is used in order to avoid S phases overlapping with the fundamental mode Rayleigh wave. Using these criteria we have ∼900 qualified events, but
more than half are eliminated after data quality control, as discussed in the next section, which leaves us
400 events (Figure 4). These events provide good azimuthal coverage, which is important for tomographic
inversions for phase velocity and azimuthal anisotropy.
2.2. Data Selection and Processing
We acquire seismic waveforms from the continuous waveform databases at the University of Alaska
Fairbanks, including the Alaska Earthquake Center and the four PASSCAL projects (section 2.1). Basic waveform processing, including instrument deconvolution, was performed using the GISMO Waveform Toolbox
for Matlab [Reyes and West, 2011]. For each event in Figure 4, a 1.5 h time window was downloaded starting from the source origin time. We used only the vertical component velocity seismograms for analyzing
Rayleigh waves, since the radial component tends to be noisier and contaminated by refracted Love waves.
Waveforms of each event were examined to eliminate bad station records such as those with missing data
segments. The waveforms were then filtered in SAC [Goldstein et al., 2003] using narrow band, zero phase
shift fourth-order Butterworth filters centered at periods of 22, 25, 29, 33, 40 46, 50, 59, 67, 77, 91, 100, 111,
125, and 143 s (Figure 5).
Data quality checking is time consuming, especially with a large number of events, but it is essential in order
to avoid propagating unmodeled data errors into the tomographic model. Records with a signal-to-noise
amplitude ratio less than 3.5 were eliminated. Further elimination of records was based on waveform similarity (cross correlation) to a reference station. Since long-period surface waves have not been used in a
comprehensive study in Alaska, we considered the station response files “untested” in some sense. We identified some suspected errors in metadata (phase response, amplitude response, and timing of clock) for the
permanent Alaska stations. Some of these errors were identified in Tape [2012] and some were corrected
in the (official) metadata. After the reliable seismograms were selected, the Rayleigh wave signals were
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Figure 4. Earthquake sources used in this study. There are 400 events with epicentral distances between 30◦ and
120◦ (large circles) of the center of the study region of Alaska, which is indicated by the star at the center of the map
projection (longitude −147◦ , latitude 62◦ ).
then isolated from other phases by windowing, tapering, and zero padding. Events with fewer than four
seismograms were discarded. With these criteria, we selected 400 events from the ∼900 candidate events.
Although the two-plane method is designed to handle erroneous instrument responses [Yang and Forsyth,
2006a], we erred on the side of caution by using the highest quality data in the inversion.
3. Methods
3.1. Surface Wave Tomography
Surface waves propagate along the interface between two different media (e.g., air and rock), with amplitude decaying exponentially with depths. There are two types of surface waves, Love waves and Rayleigh
waves. We use fundamental-mode Rayleigh waves, which are a combination of compressional (P) and shear
(SV) waves that satisfy the free-surface boundary condition. Surface waves exhibit dispersion, whereby
phase velocities are frequency dependent (Figure 5). Dispersion occurs because surface waves with different periods are sensitive to structure over different depth ranges (Figure 6), and because velocity
structures of the Earth vary with depth. Using measurements of filtered seismic waveforms from stations
at the surface, it is possible to solve a formal inverse problem to obtain the 3-D variations of Earth’s seismic
velocity structure.
Surface waves traveling large distances on the Earth have typically undergone waveform distortion due to
scattering and multipathing induced by heterogeneities. Surface wave propagation has traditionally been
approximated by a single plane wave traveling along the great circle path between the source and station.
However, this representation of the incoming waves leads to a systematic bias in the phase velocity
measurement. To alleviate this problem, Forsyth and Li [2005] developed a technique that employs a
two-plane-wave approximation and also a 2-D finite frequency sensitivity kernel representation [Zhou et al.,
2004; Yang and Forsyth, 2006b]. This technique accounts for waveform distortion, which helps improve
accuracy in the phase velocity measurement.
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25 s
29 s
33 s
40 s
46 s
50 s
59 s
67 s
79 s
91 s
100 s
111 s
125 s
10.1002/2014JB011438
The incoming surface waves are represented by
the sum of two interfering plane waves. Typically the larger one of the two waves propagates
within a few degrees of the great circle path,
while the smaller one could deviate from the
great circle path by up to 10◦ –20◦ . As discussed
in Forsyth and Li [2005], more than two plane
waves could be used to better describe the
complex wavefields, but at the cost of increasing number of parameters. However, wavefield
parameters are generally not well constrained for
global surface wave studies.
We use 2-D finite frequency sensitivity kernels
based on single-scattering (Born) approximation [Zhou et al., 2004; Yang and Forsyth, 2006b]
to account for additional wavefield perturbations caused by heterogeneities within the study
region. Due to the finite frequency effect, the
surface waves are sensitive to velocity structure
not only along the great circle path between two
seismic stations, but also in a broad region near
the great circle path. The technique provides resolution not only within the seismic array but also
in the adjacent area outside the array.
Figure 5. Example data used in this study. The waveforms
are for the vertical component, windowed Rayleigh waves
for a single event at a single station. The seismogram is
narrow-band filtered at distinct intervals between 22 s and
143 s, and two waveforms (22 s and 143 s) are excluded
due to poor signal-to-noise ratios. The filtered waveforms
reveal the basic feature of dispersion that is exploited in surface wave tomography: different periods travel at different
speeds because they are sensitive to different depth ranges of
Earth structure (e.g., Figure 6). The station is COLA.IU (latitude
67.23◦ , longitude −150.20◦ ). The event is a Mw 6.4 earthquake in the Fuji Islands region (latitude −16.54◦ , longitude
−177.52◦ ) on 31 March 2011.
The procedure of the two-plane-wave approximation and 2-D sensitivity kernel tomographic
technique falls into two primary steps. First,
phase and amplitude information of Rayleigh
waves recorded by an array of stations are
used within an inverse problem to obtain
frequency-dependent 2-D phase velocity structure and azimuthal anisotropy. Simultaneously,
wave parameters describing the incoming wavefield for each earthquake source are estimated.
Second, the 2-D maps of phase velocity and
azimuthal anisotropy are used to estimate 3-D
shear velocity and anisotropic structure by using the depth sensitivity kernels (Figure 6). This tomographic
technique has been successfully applied to a number of tectonic regions, such as the Gulf of California,
southern California, the Colorado Rocky Mountains, the East Pacific Rise, and the Tanzanian craton [Forsyth
et al., 1998; Weeraratne et al., 2003; Li et al., 2005; Yang and Forsyth, 2006a; Wang et al., 2009; Rau and
Forsyth, 2011; Wang et al., 2013]. Some of these studies have had sufficiently good data sets to resolve
anisotropic structure.
For the 2-D phase velocity inversion, the study region is parameterized by 25 by 30 nodes with 0.5◦ by 0.5◦
spacing (Figure S5). The grids extend well outside the area of station coverage, where resolution is expected
to be worse (but not zero). The seismic stations are grouped into three subregions to allow more complexity
in the incoming wavefield. By using the three subregions (instead of one), we can better account for source
effects, such as the amplitude variations due to the source radiation pattern.
Phase velocities are inverted at each grid point. We carried out two versions of the phase velocity inversion: (1) assume no anisotropy and invert for the isotropic component of velocity only at the 2-D grid points;
(2) include azimuthal anisotropy and invert for both isotropic and anisotropic components at the 2-D grid
points. In a slightly anisotropic medium, the azimuthal dependent anisotropy of Rayleigh wave can be
expressed as [e.g., Smith and Dahlen, 1973]
c(𝜃) = A0 + A1 cos 2𝜃 + A2 sin 2𝜃
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where 𝜃 is the back azimuth angle from the grid
point to the event. In the isotropic inversion, we
inverted A0 only; in the azimuthal inversion, we
inverted A0 , A1 , and A2 at each grid point.
0
−50
Depth (km)
−100
−150
−200
−250
−300
25 s
40 s
67 s
100 s
125 s
−350
−400
0
0.005
0.01
0.015
0.02
3.2. Using 1-D and 3-D Reference Models
Our tomographic velocity models are represented as perturbations from some reference
velocity model. We consider two different reference models, a 1-D layered model and a 3-D
model with more realistic crustal thickness variations. The 1-D model has the benefit that the
resultant perturbation does not reflect any
spurious 3-D crustal complexity that may be
introduced when using a 3-D reference model.
The 3-D model has the benefit that it should be
a more accurate representation of the Earth’s
crust in Alaska.
Figure 6. Fundamental mode Rayleigh wave phase velocity sensitivity kernels (c: phase velocity, 𝛽 : shear velocity). These kernels
are specific to the 1-D reference model described in section 3.2.
Longer-period waveforms are more sensitive to structures at
greater depths.
Our 1-D reference model for the crust is the
standard 1-D model for earthquake locations
in central Alaska, which is based in part on
results from the TACT active source experiment [Beaudoin et al., 1992; Ratchkovski and
Hansen, 2002a]. Below the 31 km Moho, we use the global 1-D model AK135 [Kennett et al., 1995]. This 1-D
model is applied everywhere in the model, even to the oceanic regions, where we expect it to be a poor
representation for the structure.
Our 3-D reference model is essentially the 1-D model, but stretched and compressed according to a more
realistic Moho surface that we describe in section S1. The Moho depth (below sea level) ranges from 11 km
under the Pacific plate to 48 km beneath Cook Inlet [Saltus et al., 2007]. For regions where the estimated
Moho is shallower than the Moho of the 1-D model (31 km), we compress the 1-D crustal model and stretch
the AK135 mantle model. For regions where the estimated Moho is deeper than 31 km, we stretch the 1-D
crustal model and compress the AK135 mantle model. These minor adjustments to the AK135 model are
only applied above a depth of 210 km.
3.3. Resolution Tests
Figure S6 shows two example resolution tests for the T = 40 s phase velocity maps. The resolution test provides insights into how the experimental design will influence the recovery of features with certain length
scales. Mathematically the resolution test is m = Rminput where R = G† G is the model resolution matrix, G is
the design matrix, and G† is the generalized inverse of G.
The left-hand figures in Figure S6 show minput , and the right-hand figures show m. In each resolution test
we use the same 0.5◦ nodes, but we change the length scale of the checkerboard pattern; for example, in
Figure S6 the length scales are 1.0◦ and 1.5◦ . The resolution tests show that the amplitude and checkerboard appearance of the input models are best recovered for the longer-length-scale models, as expected.
Within the region of best data coverage—near the Yakutat outline—the checkerboard pattern is adequately resolved. The reader is cautioned that these resolution tests represent best-case scenarios and do
not take into account uncertainties in data or approximations in the forward model [e.g., Lévêque et al., 1993;
Aster et al., 2012].
4. Results
We present our 3-D tomographic model in a series of cross sections of shear wave speed, such as in Figure 7
and Figure S11. Corresponding phase velocity maps are shown in Figures S7 and S8. Rayleigh waves are
sensitive to compressional and shear wave speeds. Our compressional wave speed model is simply a scaled
version of the shear wave speed model; we use a P-to-S ratio of 1.75 throughout the model.
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Figure 7. Cross sections of the shear velocity model derived using a reference model with 3-D crustal thickness (upper
left). The reference VS value (km/s) for each cross section is shown in the inset label of each subplot. The deeper cross
sections are shown with seismicity in Figure 8. The 120 km cross section is plotted in Figure 9. Figure S11 shows the same
results, but instead using a 1D layered reference model with flat Moho. The mask in each subplot is derived from the
estimated error of the phase velocity map at period 40 s (Figure S10).
We performed tomographic inversions first using a 1-D layered model with flat Moho, and then using a simple 3-D reference model with variable Moho surface (section 3.2). While there are differences between the
two resultant tomographic models (section S2), the features identified below apply to both models. Because
we expect that the 3-D reference model—notably the variable crustal thickness—is a more accurate
representation of Alaska crustal structure, we use the subsequent tomographic model in our interpretations.
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Figure 8. Deep cross sections of Figure 7 with seismicity superimposed. (The color scale is the same (±3.5%) contour line
1.0 with the VS reference value of 4.45 km/s in each cross section.) The seismicity is from the Alaska Earthquake Center
earthquake catalog, with M ≥ 2, 1990–2013. No local earthquakes were used within the tomographic inversion, so the
slab seismicity provides validation for the high-wave-speed anomaly as the top of the subducting oceanic lithosphere.
Figure 7 shows the key results of our isotropic tomographic inversion. The inversion is performed on a large
grid (Figure S5), but only the resolved regions are unmasked in the cross sections. In this section we identify
five notable large-scale features of the isotropic model.
4.1. Subducting Pacific/Yakutat Slab
The most striking feature of the tomographic model is—perhaps not surprisingly—the subducting
Pacific/Yakutat slab, which is identified as a high-wave-speed anomaly that is most apparent at depths of
80 km and greater. In Figure 8, the agreement between the slab seismicity—independent data not used in
the inversion—and the shape of the high-wave-speed anomaly provides evidence that the tomographic
inversion “is working.” The slab seismicity correlates with the leading (landward) edge of the fast anomaly,
suggesting that the majority of the fast anomaly is subcrustal oceanic lithosphere.
The cross section at 140 km depth (Figure 8) shows excellent agreement between the seismicity and tomographic model in delineating the kinked shape of the slab [e.g., Ratchkovski and Hansen, 2002b]. We also
see that the deepest seismicity (180 km, 200 km) and the fastest slab anomalies are correlated and are
west of the Yakutat block (Figure 8). It is possible that this portion of the slab is representative of “normal”
subduction, unperturbed by the subducting Yakutat crust to the east.
One region where the fast anomaly of the slab extends significantly beyond the seismic extent of the slab
is at the northern extent, near Fairbanks. As shown in Figure 9a, at 120 km depth, the fast anomaly extends
approximately 100 km north of the maximum lateral extent of the slab seismicity. This implies that the subducting Pacific/Yakutat slab extends even further from the trench than is indicated based on the Benioff
zone. Our resolution tests (Figure S6) reveal moderate recovery of features with comparable length scale at
the northern limit of the subducting slab.
4.2. Slow, Thick Crust of Southeast Alaska
The strongest low-wave-speed anomaly in our inversion is within the crust at the coast of Alaska near the
margin of the Yakutat block (Y2 in Figure 2). This is visible in the 45, 55, and 65 km cross sections in Figure 7,
based on a 3-D reference model, and also within the same cross sections from the model based on a 1-D
reference model (Figure S11). There is a similar low-wave-speed anomaly aligned with the Pacific/Yakutat
slab beneath Cook Inlet in Figure S11; however, this anomaly is not present in Figure 7, suggesting that the
thicker crust in the 3-D reference model has the effect of eliminating the anomaly. Therefore, we interpret
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Figure 9. Three cross sections of the VS model showing the tomographic contrast at the eastern edge of the Yakutat slab
(section 4.4). (a) Horizontal cross section at 120 km depth (Figure 7), plotted with the surface traces of active faults, for
reference. F = Fairbanks, A = Anchorage, MM = Mount McKinley. The VS model shows a contrast at the Y3 edge (Figure 2)
between the subducting Yakutat slab to the west and the Wrangell subduction zone upper mantle to the east (triangle).
(b) SW-NE cross section through the point (−147.5◦ , 62.5◦ ), which is used as the reference (d = 0) for the horizontal
distance, in kilometers. The dashed line at depth 120 km corresponds to the depth of the cross section in Figure 9a.
Seismicity from the Alaska Earthquake Center catalog is plotted for reference (M ≥ 2, 1990–2013); all events within 25 km
of the cross section are projected onto the profile. (c) NW-SE cross section through the point (−147.5◦ , 62.5◦ ).
that the low-wave-speed anomaly near (or within) the Yakutat block is due to crust that is thicker than
modeled (Figure S1) or to an extremely slow crust.
The recent study by Christeson et al. [2013] provides some of the strongest constraints on the Moho surface
in all of Alaska. Using land-based broadband station recordings of marine active sources, they were able to
model Pn and PmP arrivals at large distances to obtain a 3-D map of the Moho in this region. They identified an abrupt deepening step in the Moho from offshore to onshore, with Moho depths as great as 46 km.
Future surface wave tomographic inversions should use an updated Moho surface with the latest published
data, such as Christeson et al. [2013].
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4.3. Wrangell Subduction Zone
The Wrangell subduction zone is identified by a series of active volcanoes ≤25 Ma [Richter et al., 1990] overlying a weakly defined Benioff zone with sparse earthquakes down to a depth of 120 km [Stephens et al.,
1984; Page et al., 1989; Fuis et al., 2008]. This zone is marked by the segment near latitude 62.5◦ that appears
for reference in the figures in this paper (e.g., Figure 7). The current motion of the Yakutat block is to the
northwest [Fletcher and Freymueller, 2003; Elliott et al., 2010], and the crust offshore is about 30 km thick
[Christeson et al., 2010, 2013]. This means that the downgoing “Wrangell slab” (that is, the downgoing slab
beneath the Wrangell volcanoes) is subducting obliquely and is either anomalously thick like the rest of
the Yakutat slab [Rondenay et al., 2008]—a scenario partly supported by geochemical evidence [Preece
and Hart, 2004; Eberhart-Phillips et al., 2006]—or it is “normal” Pacific crust north of the Y2 boundary
(Figure 2). Denser station coverage with different imaging techniques are needed to distinguish between
the two possibilities.
Our tomographic model reveals higher-than-average wave speeds near the Wrangell subduction zone, but
the anomaly is not nearly as strong as the slab signature of the main subduction zone to the west (Figure 7).
At depths of 120 km and 140 km (Figure 7), the seismicity-derived slab segment for the Wrangell subduction zone separates relatively fast anomalies to the south (within the Wrangell slab) from slow anomalies
to the north. The correlation is even more pronounced for the inversion based on a 1-D reference model
(Figure S11).
4.4. Eastern Edge of the Pacific/Yakutat Slab
Characterizing the 3-D geometry of the subducting Pacific plate and Yakutat block is critical for understanding the dynamics of the Alaska subduction zone. There are several different published models of the slab
geometry [Zhao et al., 1995; Hayes et al., 2009; Jadamec and Billen, 2010; Fuis et al., 2008]. The primary region
of differences is between the eastern limit of Pacific/Yakutat slab seismicity and the Wrangell subduction
zone (which is absent in Hayes et al. [2009]). Lacking slab seismicity, how should these two regions be connected? Based on analysis of relocated seismicity, Fuis et al. [2008] interpreted a tear in the slab near the Y3
edge. Zhao et al. [1995, Figure 8], extrapolated the Wrangell slab to 250 km depth, then connected these
slab contours to the eastern limit of the well-defined Pacific/Yakutat contours. Jadamec and Billen [2010,
Figure S2] experimented with two different slab models, recognizing that the limited seismicity data leave
the slab geometry open for interpretation. Their first model (slabE325 ) is similar to the model of Zhao et al.
[1995]; their second model (slabE115 ) exhibits a kink-like structure that is suggested from the seismicity.
Our tomographic model shows a clear east-west contrast at the Y3 edge (Figure 2) of the subsurface Yakutat
slab. The contrast is most visible between depths of 65 and 145 km (Figure 7). Figure 9a shows the cross
section at 120 km depth, where the subducting Yakutat lithosphere exhibits markedly faster wave speeds
than the Wrangell subduction zone mantle wedge to the east. The vertical cross section in Figure 9b shows
the agreements between the deep (120 km) wave speed contrast and the sharp seismicity contrast within
the uppermost 60 km [Fuis et al., 2008]. Based on our results, we favor the kink-like slab geometry, such as
slabE115 of Jadamec and Billen [2010], rather than the smoothly connected geometry of Zhao et al. [1995].
Figure 9c shows a second cross section through the Y3 edge, with this one drawn subnormal to the
Pacific/Yakutat Benioff zone and subparallel to the Wrangell Benioff zone. In comparison with Figure 8,
Figure 9c provides a complementary view of the relationship between the slab seismicity and the fast
anomaly of the subducting oceanic lithosphere beneath the Pacific/Yakutat slab [Ferris et al., 2003; Rondenay
et al., 2008].
4.5. Western Edge of the North American Craton
We identify a high-wave-speed anomaly in the northeast region of our resolved model that is present below
depths of 75 km (Figure 7). North of latitude 65◦ , the Alaska-Canada border marks a transition from slower
wave speeds in Alaska to higher wave speeds in Canada (e.g., Figure 9). On the basis of previous results
from surface wave tomography, we interpret the high-wave-speed anomaly as the westernmost signature of the North American craton. Van der Lee and Frederiksen [2005, Plate 1] and Schaeffer and Lebedev
[2014] show a high-wave-speed anomaly in this region at the same depths, and it is connected to the larger
high-wave-speed body that characterizes the North American craton. See Polet and Anderson [1995] for a
discussion between cratonic age and the presence of high-wave-speed anomalies within the “cratonic root.”
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Figure 10. Anisotropic phase velocity map at period 40 s, which is maximally sensitive to depths between 40 and
120 km (Figure 6). The line segments indicate the fast “direction,” which is a line, not a vector. The length is in units of
km/s for the velocity perturbation. The radius of the circle represents the uncertainty of the fast direction. See Figure S15
for anisotropic phase velocity maps at other periods. (a) Dashed line shows outline of Yakutat slab from Eberhart-Phillips
et al. [2006]; thick line shows lateral extent of slab seismicity. (b) Same as Figure 10a but plotted also with the isotropic
part (in color) as well as active faults [Koehler et al., 2012]. Other phase velocity maps from the anisotropic inversion are
shown in Figure S12. See section 4.6 for details.
The “cratonic” high-wave-speed anomaly in our model is in a region that is actively deforming at the surface and is west of the active deformation front that marks the surface boundary of the craton [Leonard
et al., 2007, 2008]. The high-wave-speed tomographic signature of the craton is therefore west of the
surface signature of the undeforming stable craton.
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4.6. Anisotropy in the Alaska Subduction Zone
Thus far we have discussed results that were obtained assuming an isotropic model for VS . We performed
an inversion that allowed for azimuthal anisotropy as well, as discussed in section 4.6 and presented in
Figures S12–S15. Here we use the best-constrained phase velocity map (period 40 s) for our interpretation.
Figure 10 shows two versions of the same 40 s anisotropic phase velocity map. The eigenfunction associated
with 40 s period (Figure 6) shows that these Rayleigh waves are dominantly sensitive to structure between
depths of 40 and 120 km. The Moho surface for Alaska has maximal depths of about 50 km (Figure S1), so
the phase velocity map represents variations within the upper mantle: continental mantle north of the subduction zone and oceanic mantle south of the subduction zone. (This “upper mantle” could either be part
of the lithosphere or the asthenosphere.) The azimuthal anisotropy is represented as a line segment aligned
with the direction of fast Rayleigh wave propagation. We represent the angular uncertainties in this direction by using a circle. When the circle is larger than the line segment, we cannot resolve the anisotropy.
When the circle is much smaller than the line segment, we have the strongest evidence for anisotropy.
Here we identify three features in Figure 10, then discuss them further in section 5.
1. West of the Yakutat block, the fast direction of anisotropy is approximately trench normal and aligned in
the convergence direction. This signal spans from the Pacific plate, through the subduction zone, and into
the North America plate.
2. The fast direction of anisotropy is parallel to the margins of the Yakutat block. Referring to the labeling in
Figure 2, we see a sharp bend in the anisotropy that follows the kink between edges Y2 and Y3. There is
also alignment at Y4, where the anisotropy is sampling the mantle above the slab. There is trench-normal
alignment at Y5. The alignment is weak for the southernmost edge (Y1).
3. The fast direction of anisotropy is deflected beneath the central Alaska Range. South of the Alaska Range,
it is approximately aligned north-south, then north of the Alaska range, it deflects (and decreases) to the
east-west direction.
5. Discussion of Anisotropy
Seismic waves—Rayleigh waves, Love waves, SKS waves, S waves, and P waves—are sensitive to the components of the fourth-order elastic tensor, c = cijkl , which characterizes the form of anisotropy at any given
point within the Earth. By changing the form of c, the modeled seismic wavefield will change as well. A measurement on a seismogram at a particular station, such as a cross correlation of waveforms or a travel time
difference, can be used to infer subsurface variations in c. Each measurement, in theory, is made over a specific frequency range of the seismogram and therefore has a specific volumetric sensitivity to the cijkl within
the Earth [e.g., Sieminski et al., 2007a]. By changing the measurement in any way—a different station, a different waveform, or a different frequency range—the volumetric sensitivity to cijkl changes too. Here we will
review how these different observations are sensitive to elastic anisotropy, and we will next discuss them in
the context of previous work in Alaska. For extended discussion of seismological observations and sensitivities, the reader is referred to Montagner and Nataf [1986], Montagner et al. [2000], Chen and Tromp [2007],
Karato et al. [2008], Long and Silver [2009], and Becker et al. [2012].
Rayleigh waves comprise P and SV motion and propagate in the horizontal direction. One possible scenario is that azimuthal anisotropy represents hexagonal crystal symmetry with the symmetry axis in the
horizontal plane. Vertically polarized shear waves (SV), such as those that occur within Rayleigh waves, will
travel fastest in one azimuthal direction. The depth sensitivity of the Rayleigh wave depends on the frequency content, with longer period sensitive to greater depths (Figure 6). The typical azimuthal anisotropy
phase velocity map (Figure 10) is therefore tied, to first approximation, to a depth range via Figure 6 and a
particular simplified version of the elastic tensor c. Lateral resolution is controlled by the density of source
station path coverage and by the frequency content of the measurements: shorter periods have narrower
finite-frequency lateral sensitivities (section 3.1).
SKS waves start as S waves, convert to P waves in the liquid outer core, then emerge as shear waves (SV),
with the (theoretical) particle displacement in a direction that is perpendicular to the ray path and within
the source station plane. For a given station, the teleseismic earthquakes are selected such that the SKS
waves arrive at steep incident angles 10◦ < 𝜃 < 15◦ . Under the same structural assumption of azimuthal
anisotropy, the shear wave will be split into a slow- and fast-propagating wave. SKS splitting measurements
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are typically made at periods between 3 and 10 s, corresponding to finite-frequency sensitivities that are
∼100 km laterally and several hundred kilometers vertically [Sieminski et al., 2008; Long et al., 2008]. In
comparison to surface wave estimates of azimuthal anisotropy, SKS splitting measurements provide better
lateral resolution but poor vertical resolution.
S waves from local earthquakes, notably deep earthquakes within the subducting slab, can also be used
to identify anisotropy. Splitting measurements from local earthquakes are made with incident angles less
than about 35◦ [e.g., Paulssen, 2004]. Measurements from local slab earthquakes tend to be made at shorter
periods—from 0.1 Hz to 8 Hz in three examples [Nakajima and Hasegawa, 2004; Long and van der Hilst,
2006; Hammond et al., 2010]—and, therefore, the sensitivity kernels are narrower. (The shorter path lengths
also lead to narrower kernels than the longer paths of SKS waves.) Long and Silver [2009] caution that local
splitting at high frequencies could arise from crustal anisotropy, especially considering that the splitting
direction tends to emphasize the anisotropic structure closest to the receiver.
P waves are also influenced by anisotropic structure. P wave anisotropy is typically inferred from travel
time measurements from local earthquakes, the same measurements used in isotropic inversions [e.g.,
Eberhart-Phillips and Henderson, 2004; Eberhart-Phillips and Reyners, 2009; Wang and Zhao, 2008]. Examples
of sensitivity kernels for global P and S waves are presented in Sieminski et al. [2007b]. In local earthquake
studies the P arrival times are picked from raw or very high frequency waveforms, and therefore, the sensitivities are implicitly represented by rays. In the tomographic inversion, by allowing the P velocity to vary
at each node as a function of direction, it is possible to use P wave arrival times to estimate a 3-D model of
VP anisotropy.
5.1. Anisotropy in Alaska
There are many different observational and modeling constraints that must be considered when interpreting anisotropic structure within the Alaska subduction zone:
1. VS azimuthal anisotropy derived from Rayleigh waves [this study; Figure 10]
2. Shear wave splitting observations from SKS waves from teleseismic earthquakes [Christensen and Abers,
2010; Hanna and Long, 2012]
3. Shear wave splitting observations from S waves from local slab earthquakes [Wiemer et al., 1999]
4. VP anisotropy derived from travel times of P waves from local slab earthquakes [Tian and Zhao, 2012]
5. Numerical modeling results of predicted flow fields [Jadamec and Billen, 2010]
6. Measured anisotropy from exhumed upper mantle rocks [Mehl et al., 2003]
We are confident that some anisotropy is present in the Alaska subduction zone. The challenge lies in
determining what kinds of anisotropy are present and how 3-D variations in anisotropy are distributed
throughout the crust and mantle of the subduction zone. (By “subduction zone,” we mean the upper
∼800 km of the region spanning from the incoming oceanic Pacific plate to the downgoing slab and into
the overriding continental North America plate.) Our discussion is in terms of intrinsic anisotropy, but we
acknowledge the possibility of apparent anisotropy caused by heterogeneous structures [e.g., Backus, 1962;
Fichtner et al., 2013].
We provide three new observations of anisotropy in Alaska, listed in section 4.6. We will discuss these in the
context of previous studies, with emphasis on SKS splitting results. Perttu et al. [2014] presented SKS splitting
measurements from three seismic deployments covering a north-south swath across the entire state. These
observations reveal a strikingly simple overall pattern in Alaska: south of the 70 km subduction interface
contour of Syracuse and Abers [2006], the fast splitting directions are NW-SE; north of the contour they are
NE-SW. Figure 11b shows a zoom-in on the abrupt transition in the SKS splitting measurements. Our fast
directions modeled from long-period Rayleigh waves are plotted for comparison in Figure 11c and show a
similar deflection from N-S alignment in the south to E-W alignment to the north.
A second area of agreement between our 40 s anisotropy map and the splitting results is in the
trench-normal fast directions beneath the Kenai peninsula, west of the Yakutat crust. Perttu et al. [2014] favor
an interpretation of asthenospheric flow, but they also acknowledge the possibility of anisotropy within
the subducting oceanic plate [Song and Kawakatsu, 2012, 2013]. These fast directions are approximately
aligned with the convergence direction and also with the absolute plate motion direction for the Pacific
plate (Figure 11a). (As noted in Perttu et al. [2014], the trench-parallel splitting at stations closer to the trench
requires a different interpretation.)
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Figure 11. SKS shear wave splitting measurements in Alaska [Christensen and Abers, 2010; Perttu et al., 2014]. (a) Fast
direction is plotted as black segments, with the length proportional to the delay time between the two orthogonal
shear waves. Delay times range from 1 to 2 s. PA: Pacific plate; YK: Yakutat block; NA: North America plate. Red triangles
represent active volcanoes. Arrows show plate motion vectors, which are plotted at uniformly spaced locations, with four
vectors at each location. All vectors are based on the plate model of Bird [2003] but with one of four different reference
frames: black = fixed North America plate; white = no-net-rotation [DeMets et al., 1994]; red = hotspot of Gripp and
Gordon [2002]; cyan = hotspot of Morgan and Morgan [2007]. (b) Zoom-in on the key transitional region beneath the
Alaska Range. Colored contours of the subduction interface are from the global 3-D subduction interface model Slab
1.0 [Hayes et al., 2009]. The interface in this region is based on relocated slab seismicity from Fuis et al. [2008]. The lateral
transition in the splitting observations occurs at a slab contour. The depth of the contour depends on the choice of slab
model: it is 90 km for Slab 1.0 (plotted here) or 70 km for Syracuse and Abers [2006]. (c) Zoom-in on Figure 10 to show the
surface wave anisotropy results for comparison.
There are some notable differences between our results and the splitting observations. West of the Yakutat
crust, there is a NW-SE trend in the fast directions that spans from the Pacific plate and into the back-arc
region (section 4.6). Although there are fewer stations in this region for splitting observations, the limited
results substantiate the deflection that is identified to the northeast beneath the subducting Yakutat crust
[Perttu et al., 2014]. Our results suggest that the Yakutat crust is playing a primary role on changes in fast
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direction, since we do not see any deflection west of the Yakutat crust. This differs from the simpler interpretation, based on the SKS splitting observations, that anisotropy abruptly changes due to the presence of a
mantle wedge near the 70 km slab contour [Perttu et al., 2014].
It is likely that a joint inversion of the anisotropic phase velocity maps with the SKS splitting observations
[e.g., Lin et al., 2011] would lead to a 3-D anisotropy model that predicts fast directions that are consistent
for both the Rayleigh wave data and the SKS observations. Furthermore, improving the poor station coverage in the back-arc region of southwestern Alaska will be important for resolving the apparent discrepancy
between results from SKS splitting and surface wave anisotropy.
6. Conclusion
We present a 3-D shear wave velocity model for the Alaska subduction zone (Figures 7–9) that provides a
valuable, large-scale structural framework for future studies in this complex zone of subduction, collision,
and broad plate deformation. Major tomographic anomalies in our model include the following: (1) high
wave speed of the subducting Pacific/Yakutat slab (section 4.1), (2) slow, thick crust in southeast Alaska
within the Yakutat domain (section 4.2), (3) lower wave speeds associated with the Wrangell subduction
zone (section 4.3), (4) strong contrast at the eastern edge of the Pacific/Yakutat slab between depths of
65 km and 120 km (section 4.4), and (5) high wave speed in Canada, interpreted as the westernmost signature of the North American craton. By using teleseismic surface waves, we are able to cover a broad region,
including outside the array of stations, albeit at a lower resolution than local tomography body wave studies
[e.g., Eberhart-Phillips et al., 2006].
The 3-D isotropic shear wave structural model is complemented by a set of anisotropic phase velocity maps,
the best resolved of which is derived from measurements of 40 s surface waves (Figure 10). These maps
provide a valuable perspective on anisotropy that is complementary to previous studies on SKS splitting
[Christensen and Abers, 2010; Hanna and Long, 2012; Song and Kawakatsu, 2013; Perttu et al., 2014].
Our results provide additional evidence for the dominant influence of the Yakutat collision on active tectonics in Alaska (e.g., section 1.1). Subduction in southern Alaska has persisted since at least 160 Ma [Fisher
and Magoon, 1978; Rioux et al., 2010]. The relatively recent arrival of the Yakutat block at 24 Ma [Benowitz
et al., 2011] has left an indelible mark in the form of flat slab subduction, an extinguished volcanic arc, and
broad intraplate deformation, among others. The isotropic model shows a sharp contrast at the edge of
the Yakutat slab (Figure 9), as well as slow wave speeds near the coast, where the collision is active. The
anisotropic results show a clear transition from NW-SE fast directions west of the Yakutat slab to a more variable pattern within the region of the (inferred) Yakutat slab (Figure 10). Furthermore, the fast directions are
parallel to almost all sides of the Yakutat slab, which favors an interpretation based on flow rather than on
relic structures.
Understanding the anisotropic structure of Alaska will require new seismic data, as well as more sophisticated modeling approaches. The geodynamic models of Jadamec and Billen [2010] and Jadamec et al.
[2013] show the great challenge involved in modeling 3-D processes at the edge of a subduction zone with
such complex geometry. Jadamec and Billen [2010] were able to test the influence of slab geometry (notably
the Wrangell slab) and viscosity on the resulting flow field, whose splitting predictions produced modest
agreement with SKS splitting observations. A similarly sophisticated approach will be required to model
seismic anisotropy in this region. Numerous studies have identified seismic anisotropy from different seismic waveform measurements (section 5), but most efforts are unable to convincingly isolate the region of
sensitivity and the style of anisotropy giving rise to the measured waveforms. In the future we hope to use
3-D seismic wavefield simulations to test various types of anisotropy within the context of a realistic 3-D
isotropic reference model that includes the subducting Pacific/Yakutat slab.
Our tomographic inversion exploited several dense, temporary seismic deployments (Figure 3). The
expected deployment of the Transportable Array in Alaska will fill in large gaps in coverage, such as in southwestern Alaska and western Canada. Future tomographic inversions with teleseismic surface waves should
be able to improve upon the results in our study, just as a statewide map of SKS splitting will improve upon
the results of Perttu et al. [2014]. However, some features that we have discussed, such as the details in
the vicinity of the Wrangell subduction zone, require both denser station spacing as well as imaging with
higher-frequency waves, either from teleseismic body waves or from local slab or crustal earthquakes.
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Acknowledgments
We used seismic waveforms from all
permanently operating stations in
Alaska, in addition to several temporary seismic experiments (section 2.1).
We thank all those who were responsible for installing and maintaining
these stations. All waveforms were
obtained from the Alaska Earthquake
Center data archive and are available
at the IRIS Data Management Center. We thank Doug Christensen and
Geoff Abers for providing their SKS
splitting observations in Figure 11
prior to their publication [Perttu et al.,
2014]. Y. Wang was supported by a
postdoctoral fellowship at the Geophysical Institute, University of Alaska
Fairbanks. C. Tape was supported by
a grant from the National Science
Foundation (EAR-1215959). We thank
Andreas Fichtner and an anonymous
reviewer for their constructive reviews.
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